Ore Geology Reviews 43 (2011) 333–346

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Ore Geology Reviews

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Gushan magnetite–apatite deposit in the Ningwu basin, Lower Yangtze River Valley,SE China: Hydrothermal or Kiruna-type?

Tong Hou a, Zhaochong Zhang a,⁎, Timothy Kusky b

a State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Beijing, 100083, Chinab State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan, 430074, China

⁎ Corresponding author. Tel.: +86 10 82322195; fax:E-mail address: zczhang@cugb.edu.cn (Z. Zhang).

0169-1368/$ – see front matter © 2011 Elsevier B.V. Alldoi:10.1016/j.oregeorev.2011.09.014

a b s t r a c t

a r t i c l e i n f o

Article history:Received 12 January 2011Received in revised form 15 September 2011Accepted 16 September 2011Available online 21 September 2011

Keywords:MetallogenesisKiruna-type depositsGushanLower Yangtze River ValleyLiquid immiscibility

The Gushan deposit is one of the typical magnetite–apatite deposits associated with dioritic porphyries in theLower Yangtze River Valley belt of the eastern Yangtze craton. The origin of this deposit is still uncertain andremains a controversial issue. Divergent opinions are centered on whether the iron deposits are magmatic orhydrothermal in origin. However, our field observations and mineralogical studies, combined with previouspublished petrological and geochemical features strongly suggest that the main ore bodies in the Gushanmagnetite–apatite deposit are magmatic. Specific evidence includes the existence of gas bubbles, tubes,and miarolitic and amygdaloidal structures, melt flow banding structure and the presence of “ore breccia”.New electron microprobe analyses of the pyroxene phenocrysts of the dioritic porphyry genetically associat-ed with the Gushan magnetite–apatite deposit show that the Fe contents in the evolving magma dramaticallydecrease, and then gradually increase. Because there is no evidence of mafic magma recharge, this scenario(decreasing Fe) could be plausibly interpreted by Fe-rich melts separated from Fe-poor silicate melts, i.e., liq-uid immiscibility was triggered by minor addition of phosphorus by crustal contamination. The occurrence ofmassive iron ore bodies can be satisfactorily explained by the immiscible Fe-rich melt with enormous volatilecontents was driven to the top of the magma chamber due to the low density. The hot and volatile-rich ironore magma was injected along fractures and spaces between the dioritic intrusions and wall-rocks, and led toan explosion near the surface, resulting in the immediate fragmentation of the roof of the intrusion and wall-rocks, forming brecciated ores. Moreover, other types of ores can be considered as a result of post-magmatichydrothermal activities. Our proposed metallogenic model involving the Kiruna-type mineralization is con-sistent with the observed phenomenon in the Gushan deposit.

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1. Introduction

Magnetite–apatite deposits, also known as Kiruna-type deposits,occur in a number of locations in theworld, range in age from Proterozo-ic to Holocene, and are associatedwith volcanic rocks or sub-volcanic in-trusions (Nyström and Henriquez, 1994). Controversy persists regardingthegenesis of enigmatic Kiruna-type deposits dominated by sulfide-poormineral assemblages of low-Ti magnetite–fluorapatite–actinolite, andrange in size from large bodies containing many hundreds of millionsof tons of high-grade iron ore, to small veins and veinlets (Hildebrand,1986; Nyström and Henriquez, 1994). The Kiruna-type deposits havebeen interpreted to have an exhalative-synsedimentary (Aftabi et al.,2009; Parák, 1975), or alternatively epigenetic–hydrothermal origin(e.g., Bookstrom, 1977; Gleason et al., 2000; Hildebrand, 1986; Jami etal., 2007, 2009; Sillitoe and Burrows, 2002), or a magmatic origin (liquidimmiscibility) (e.g., Förster and Jafarzadeh, 1994; Frietsch, 1978;

Naslund et al., 2000, 2002; Nyström and Henriquez, 1994). There iswidespread petrologic evidence for the formation of such liquids andfor their immiscible relationships with felsic melts (the type (I) immisci-bility of Naslund, 1983), particularly in the tholeiitic basalt environmentinwhich residual liquids are commonly enriched in iron owing to under-saturation in ferroan spinel during early crystallization stages (Philpotts,1967; Roedder, 1979). Current understanding tends to group the ‘Kiru-na-type’ deposits as an end-member of the hydrothermal Iron Oxide–Copper–Gold ‘IOCG’ clan (Hitzman, 2000; Hitzman et al., 1992). This issupported by similarity of tectonic setting, abundance of early-stagemagnetite, occurrence of minor late stage pyrite, chalcopyrite–gold–REE in or nearmassivemagnetite deposits, and certain shared post-mag-matic and gangue minerals, especially actinolite and apatite. Barton andJohnson (1996, 2000) suggest that evaporites were a source of Cl, thetypical Na alteration, and the high oxidation state of such deposits,which involves a process of large-scale basinal brine circulation drivenby intrusion buried. Obviously, this inference needs to be tested byfield and petrographic observation.

The Ningwu volcanic basin is located in the east part of the LowerYangtze River Valley (LYRV), an important Cu–Au–Fe–S ore belt

334 T. Hou et al. / Ore Geology Reviews 43 (2011) 333–346

associated withMesozoic magmatic rocks in SE China. A large numberof magnetite–apatite deposits occur in the Ningwu basin occur(Fig. 1), and are mainly characterized by a magnetite–apatite–actino-lite assemblage that can be well compared with the Kiruna-type mag-netite–apatite deposits (Yu et al., this issue; Jiang et al., 2006; Song etal., 1981; Yu and Mao, 2002), and were named “porphyry-type irondeposits” by Chinese researchers in 1970s (Ningwu Research Group,

Fig. 1. Geological sketch map of the Ningwu basin (simplified from the Ningwu Research GrGushan, Baixiangshan, Hemushan the data are from Fan et al. (2010), for Niangniangshan f

1978). These deposits have produced over 2000 million tons (Mt)of ore containing an average of 45 wt.% FeO, and currently stillhave 2340 million tons of ore (Gao and Zhao, 2008). Recently, twolarge Kiruna-type deposits, Yangzhuang and Nihe, with N100 Mt ofhigh grade ores have been discovered in the Ningwu basin and adja-cent Luzong basin, respectively (e.g., Lin et al., 2010). Previous inves-tigations have suggested that the dioritic magmas probably provided

oup, 1978). The age data for Jishan is from Hou and Yuan (2010); for Taocun, Washan,rom Yan et al. (2009) and for Longwangshan from Zhang et al. (2003).

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a heat source as well as a potential fluid, metal, and sulfur source, forthe ca. 130 Ma iron oxide mineralization (e.g., Hou and Yuan, 2010;Hu and Jiang, 2010).

Although these magnetite–apatite deposits in the Ningwu basinhave attracted a substantial number of petrologic and geochemicalstudies (e.g., Chang et al., 1991; Ningwu Research Group, 1978; Panand Dong, 1999; Xu, 1990; Zhai et al., 1996), their metallogenesisand the genetic relationship with associated igneous rocks are stillpoorly understood, and therefore the origin of magnetite–apatite de-posits in this region is still debated. Accordingly, a hydrothermal ori-gin was first proposed by Geological Survey of Industrial Ministry(1935), and was then adopted by Lin et al. (1983), Institute of Geo-chemistry of Chinese Academy of Sciences (1987), Lu et al. (1990)and Gu and Ruan (1988), although they found that some evidencefor the stock of magnetite magma and formation through fumarolicactivity in these deposits merits serious consideration. Thus, there isdoubt as to whether features described as being typically magmaticare indeed so, as opposed to being either a hydrothermal re-arrange-ment of magmatic features or a primary hydrothermal feature. Be-sides, those authors who favored the magmatic model for theporphyry-type mineralization in China only tried to provide field ev-idence for the involvement of iron ore magma (e.g., Zhai et al., 1996),but many important issues remain to be resolved, e.g., if they are pri-mary magmatic ore, whether the liquid immiscibility has been in-volved in the mineralization and how liquid immiscibility wastriggered, and furthermore, why almost all the main ore bodies inthe region occur in the roof zones (apical parts) of the intrusions in-stead of staying behind in the magma chamber which is imaginableas a result of their relatively high density in the magmatic plumbingsystems (e.g., Gibson, 2002; Veksler et al., 2006).

The Gushan magnetite–apatite deposit, typical of the Lower YangtzeRiver Valley, is located in the southern part of the Ningwu basin. The ori-gin of this deposit is still uncertain and remains a controversial issue. Di-vergent opinions are centered onwhether the iron deposits aremagmaticor hydrothermal in origin. Therefore, it provides a rare opportunity to ex-amine the variousmetallogeneticmodels, and such a re-examination is inturn essential to elucidate their origin and can greatly improve the prev-alent model of the world-wide Kiruna-type deposits (Frietsch, 1991;Frietsch and Perdahl, 1995; Martinsson, 1997; Martinsson and Weihed,1999; Nyström and Henriquez, 1994; Ripa, 1988). In this paper, we sys-tematically study the geology and geochemistry of the Gushan depositsin order to provide some crucial clues and shed new light on the genesisof shallow Kiruna-type mineralization in particular, which is a necessaryprerequisite for an improved and robust genetic model of the relation-ships between dioritic porphyries and associated Fe mineralization.

2. Regional geology

TheNingwubasin is situatednear the northernmargin of the YangtzeBlock to the east of Tangcheng–Lujiang fault (Pan andDong, 1999). It ex-tends for approximately 80 km from the Meishan orefield in JiangsuProvince in the north to the Zhonggu orefield of Anhui Province in thesouth, forming a north-northeast trending rhomboid-shaped faulted vol-canic basin (Fig. 1), as one of the seven ore clusters in Lower YangtzeRiver Valley (LYRV). The basement rocks in the Yangtze Block includeamphibolite and granulite facies biotite-hornblende gneisses, tonalites,trondhjemites, granodiorites and supracrustal rocks, exhibiting perva-sive migmatization (Zhai et al., 1992). Zircon U–Pb and whole-rockSm–Nd geochronological data reveal that these basement rocks arePalaeo-Proterozoic to Archean in age, 1895 to 2900 Ma (Chang et al.,1991). The basement is overlain by a 2000-m-thick Palaeo- to Neo-Proterozoic (990 to 1850 Ma,Chang et al., 1991) volcano-sedimentarysuite of calc-alkaline basalts, rhyolitic rocks, and shallowmarine carbon-ate and clastic sedimentary rocks that have been moderately metamor-phosed to schists and gneisses (Pan and Dong, 1999). A recentbiostratigraphic and stratigraphic investigation has suggested that,

starting in the Cambrian, thick (~1 km) carbonate and clastic sequenceswere deposited in the Palaeotethys ocean, and a large number oforganic-rich black shales and chert nodules as well as phosphorouslayers and nodules were formed in response to several anoxic eventsduring the Palaeozoic (Lü et al., 2004). There are many P2O5-rich bedsvisible along theYangtze River inAnhui Province in Cambrian to Permianstrata which contain some phosphorus-rich layers such as themarlstonerocks of the lower-middle Jurassic Xiangshan Group with average P2O5

concentrations of up to 6.93 wt.% (Zhao, 1993a,b), and the PermianGufeng Formation which consists predominantly of gray chert andblack siliceous shales with chert nodules intercalated with 11 to107 m-thick phosphorous-rich layers (Chang et al., 1991). The Ningwuvolcanic basin is largely covered by continental volcanic rocks intrudedby cogenetic subvolcanic and plutonic rocks. Sedimentary rocks of theUpper Triassic Huangmaqing Formation, lower-middle Jurassic Xiang-shan Group and upper Jurassic Xihengshan Formation constitute thebasement of the volcanic basin, and are also exposed in the surroundingarea. The volcanic rocks in the Ningwu basin have basaltic and trachyan-desitic affinities. They are enriched in sodium and have a trend from fair-ly basic to intermediate and alkalic. The average chemical composition isof a basalt-trachyandesite, (55.93 wt.% SiO2, 4.40 wt.% Na2O, 3.17 wt.%K2O; Ningwu Research Group, 1978). Four volcanic formations havebeen recorded, including the Longwangshan Formation (about20%), Dawangshan Formation (about 75%), Gushan Formation, andNiangniangshan Formation (about 5%), from bottom to top. The vol-canic rocks of the Niangniangshan Fm. are exposed only around theNiangniangshan area in the middle part of the Ningwu basin. TheSHRIMP zircon U–Pb ages of Longwangshan Fm., Dawangshan Fm.and Niangniangshan Fm. volcanic sequence is 131±4, 127±3 Maand 130.6±1.1 Ma, respectively (Yan et al., 2009; Zhang et al.,2003), indicating an early Cretaceous age. The Niangniangshan For-mation is characterized by the presence of peralkaline rocks, con-sisting predominantly of nosean phonolite and minor trachydaciteas well as minor corresponding subvolcanic rocks, e.g., pseudo-leucite porphyry, whereas the Longwangshan, Dawangshan andGushan Formations are characterized by subalkaline rocks, consist-ing of (basaltic) trachyandestic and andesitic rocks. Mineralizationin the basin is chiefly related to subvolcanic rocks (gabbro–diorite–gabbro–dioritic porphyries) that formed in the late stage of deposi-tion of the Dawangshan Formation. Furthermore, the spatial distri-bution of these subvolcanic rocks is controlled by the intersectionsof two groups of faults trending NNE and WNW, respectively. Thevolcanic edifices as well as subvolcanic intrusions and plutons areoften found at the intersections of faults. It has been shown by theNingwu Research Group (1978) that the magma was of basalt-andesitic composition enriched in alkalis and volatiles (1.51 wt.%H2O and much F and Cl, leading to formation of a large amount of ac-cessory apatite and pegmatitic fluorite–apatite–diopside veins).

The magnetite–apatite deposits in the Ningwu basin are mainlyclustered in the northern Meishan, central Washan and southernZhonggu orefields from north to south (Fig. 1). The host rocks are ex-clusively intrusive with minor sedimentary beds. The intrusions aretypically subvolcanic dioritic porphyries, and intruded the sedimenta-ry and volcanic rocks as stocks. The porphyry iron mineralization oc-curs in massive and disseminated ores in apical parts of intrusions(cupolas) of the porphyry stocks, and in lodes in the volcanic rocks,and are typically accompanied by albite alteration with the assem-blage albite-(or marialite)–actinolite (or diopside)–apatite. Based onthe geophysical data and deep drilling information, from ~−1500 mupwards to ~−800 m, there is possibly a large mafic batholith atdepth which probably links together all the dioritic porphyries inthe Ningwu basin (Chang et al., 1991). Moreover, extensive geochem-ical studies on the dioritic porphyries which host iron ores suggestthat they were derived from an incompatible element-enriched litho-spheric mantle, followed by fractional crystallization and crustal con-tamination (AFC; Hou et al., 2010; Xing, 1998).

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3. Geology of the Gushan magnetite–apatite deposit

The Gushan dioritic stock has an outcrop area of ~0.3 km2 and is ahypabyssal intrusive stock with the shape of a half ellipsoid. The stockintruded sandstone and shale of the Huangmaqing Fm and quartzsandstone and arkosic sandstone of the Xiangshan Group (Fig. 2).The dioritic porphyry consists of clinopyroxene (~20%) and plagio-clase (~60%) phenocrysts with a fine grained groundmass, whichhas intersertal to intergranular texture, and consists predominantlyof plagioclase and augite with minor accessory anhedral iron oxides,fluorapatite and titanite (Fig. 3a). The clinopyroxene content in-creases and the size of the grains in the porphyries increase withdepth. Slightly altered labradorite rimmed by fresh albite is common-ly present. However, some albites are present in the groundmass, andare eutectic and fine-grained, suggesting a magmatic origin. Extensivecarbonate and kaolinitic alteration is associated with large-scale ironmineralization.

The iron ore bodies have been recognized at the contact zone be-tween the dioritic stock and the sedimentary country rocks. TheGushan deposit is characterized by variable ore body shapes forminga ring in the east and an irregular shape in the west. There are alsoNE-trending veined ore bodies at the southwest side of the deposit.As a whole, the ore bodies constitute a dome shape which is probablycontrolled by the shape of the dioritic porphyry intrusion and fracturesystem. The emplacement of the intrusion led to the ring shaped orebodies, and the intense fracturing resulted in the regular linear andveined ore bodies. Four types of ores have been recognized. Themost important types are massive ores and brecciated ores whichmake up 90% of the ores by volume. Almost all massive ores andpart of the brecciated ores have a high Fe grade of ~45 wt.%. Stock-work, disseminated, banded and skeleton ores are also common, butare not as important as the two above types economically. The mas-sive ore makes up 40 vol.% of the deposit. The massive ores are non-porous and consist predominantly of magnetite and martite, withminor accessory minerals such as apatite, quartz, and a few titanitegrains. In massive ores, there are tubes, and miarolitic and amygdaloi-dal structures which are filled by quartz and calcite crystals. Part ofthe massive ores occur as schlierenlike veins or dikes typically form-ing a complex network, the so-called “ore breccia”. The ore is con-formable to unconformable with the structures in the country rocks.At the lower part of the ore bodies, gas bubbles, tubes, and amygdalesare absent in the massive ores, but melt flow structures (Fig. 3b) arecommonly present instead, which are also common in the Kiruna de-posit in Sweden (e.g., Frietsch and Perdahl, 1995; Martinsson, 1997;Martinsson and Weihed, 1999; Nyström and Henriquez, 1994). Thedecrepitation temperatures of inclusions in magnetite and martiterange from 350 to 1040 °C (Li and Xie, 1984), and their temperature

Fig. 2. (a) Cross-section, and (b) geological map showing the relationships between dioriticThe straight lines in the geological maps indicate the orientation of the profile.

of homogenisation (Th) shows a peak at 1000 °C. The banded oresare always found in the deepest part of the ore bodies (Zhai et al.,1992).

4. Analytical methods

Electron microprobe analyses were determined for some pheno-crysts using an EMPA-1600 Superprobe at the State Key Laboratoryof Geological Process and Mineral Resources of the China Universityof Geosciences, Beijing. Operating conditions were set at 15 kV at10 nA beam current. Natural minerals and synthetic pure oxidesfrom SPI Company of America were used as standards. For pyroxene,the calibration standards used were hornblende (for Si, Ti, Al, Fe, Ca,Mg, Na and K), fayalite (for Mn) and Cr2O3 (for Cr). For plagioclase,the standards used were hornblende (for Si, Ti, Al, Fe, Ca and Mg), al-bite (for Na), orthoclase (for K) and fayalite (for Mn). Precision is bet-ter than 1 wt.% for element oxides. Analyses of fresh clinopyroxenesand plagioclase phenocrysts from representative Gushan dioritic por-phyry samples are reported in Tables 1 and 3, respectively, and forclinopyroxenes and plagioclase in the groundmass, the results arelisted in Tables 2 and 4, respectively.

5. Results

Analyses of phenocrysts, i.e., clinopyroxene and plagioclase, andclinopyroxene and alkali feldspar in groundmass from Gushan pro-vide more constraints on changing magmatic conditions during crys-tallization (Tables 1, 2, 3 and 4). All analyzed plagioclase phenocrystscan be classified as labradorite (An53~70), and generally, show rela-tively simple compositional and textural variations i.e., from core,mantle to rim, the An content gradually decreases, whereas the alkalifeldspars in the groundmass are mainly Na-rich orthoclase andanorthoclase and do not show zoning. All analyzed clinopyroxenesbelong predominantly to diopiside. The clinopyroxene phenocrystsshow concentric zoning with oscillatory chemical variations, whereasthe clinopyroxenes in the groundmass are homogeneous and do notshow zoning and their compositions are close to the marginal partsof the clinopyroxene phenocrysts. The principal difference betweencore and rim compositions in the clinopyroxene phenocrysts is inMgO and FeO contents, with concentrations varying by as much as6 wt.%. Particularly, from core to rim, the FeO content (~5 wt.%) in-creases in the central parts reaching up to a peak at ~9 wt.%, andthen decreases sharply to a minimum ~5 wt.%, followed by graduallyincrease to 10 wt.% at the outer rim. In contrast, the MgO content de-creases from the central parts (N16 wt.%) reaching down to a mini-mum at ~13 wt.%, and then increases sharply to a maximum

porphyry, wall-rocks and iron ore bodies at Gushan, simplified from Zhai et al. (1992).

Fig. 3. (a) Porphyritic-like texture displayed in the Gushan dioritic porphyry containing large phenocrysts of clinopyroxene (Cpx) and plagioclase (Pl) in a groundmass of fine-texturedcrystals, which consist predominantly of plagioclase and augite withminor accessory anhedral iron oxides, fluorapatite and titanite; (b) magmatic flow banding structure inmassive ironore, Gushan; (c) Gas bubbles in massive ore; (d) brecciated ore consisting of angular brecciated fragments of the country rocks and the dioritic porphyries cemented by Fe-bearingminerals.

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~16 wt.%, followed by a gradually decrease to ~12 wt.% at the outerrim.

6. Discussion

6.1. Ore genesis: hydrothermal or magmatic?

There has been considerable debate over whether Kiruna-typeores have a magmatic origin with the debate summarized by Hitzmanet al. (1992), Nyström and Henriquez (1994) and Bookstrom (1995).In terms of mineralogy and host rock lithologies, the Gushan depositscould be classified as Kiruna-type. Thus we need to consider whetherthey may be primary magmatic ores or not. There is considerable con-fusion as to what may define a magmatic deposit. Many features de-scribed for intrusive Fe-rich deposits (narrow breccia pipes,incorporation of brecciated xenoliths of country rock, needle-likegrowths of actinolite after diopside and apatite perpendicular tovein walls or breccia fragment surfaces; see for instance, Zhai et al.,1992, 1996) may equally well fit hydrothermal deposits located with-in open fractures, the early stages of which are marked by hydrother-mal brecciation.

Based on field observations in the Gushan deposit, the formationof iron ores was indeed accompanied by alteration of the host rocks,although the textural characteristics of these precursor rocks arelargely preserved. Peripheral to the mineralized zone, the countryrock is locally replaced by silicate, kaolinite and carbonate whichform veins or irregular masses, often adjacent to the ore, and restrict-ed in the contact zone between the dioritic porphyries and wall rocks.Additionally, the dioritic porphyries are characterized by a grayish-green color associated with the development of secondary chlorite,kaolinite and carbonate replacement of primary clinopyroxene andplagioclase phenocrysts respectively, as well as small amounts ofquartz, barite and fluorite. The alteration of magnetite to hematite isconnected with this process. However, the conclusion that theGushan deposit is hydrothermal in origin drawn mainly on basis of

these replacement features is still quite doubted. The following linesof evidence compel us to argue against the hydrothermal geneticmodel: (1) the existence of gas bubbles, tubes, and miarolitic andamygdaloidal structures which are filled by quartz and calcite crystalsin massive ores provide evidence of crystallization from volatile-richmelts (Fig. 3c); and (2) the sharp contact between schlierenlike mas-sive ore bodies and the dioritic porphyry, irrespective of whether theore occurs as large bodies or as small veins, with dike and pluglike ex-tensions outwards into the country rocks; (3) the presence of “orebreccia” which mainly includes fresh fragments cemented by Fe-rich minerals (Fig. 3d) and the flow banding structure in massiveore are suggested to be diagnostic of a magmatic origin; (4) the pres-ence of melt inclusions in the apatites that were trapped at magmatictemperatures (e.g., Li and Xie, 1984); (5) fine-grained magnetite(martite)–apatite mineralogy; (6) generally weak alteration of thesilicate rocks above the ore; (8) the δ18O compositions of hematitein massive ores with flow structure are +0.44‰ (Li and Xie, 1984).Conclusively, even though hydrothermal fluids contribute a greatdeal to the ore formation, the initial massive ores are undoubtedlymagmatic in origin. On the other hand, some of the field and geo-chemical relationships are readily modified by post-magmatic hydro-thermal fluids.

In addition, in the TiO2 vs. V2O5 diagram (Fig. 4), the V and Ti con-tents in the magnetite of the Gushan deposit display a transition fromthose of Kiruna and El Laco deposits to the hydrothermal field, sug-gesting there is evidence for both magmatic and hydrothermal ores,with early high temperature ores being the product of magmatic seg-regation and later, lower temperature ores having a hydrothermal or-igin. The magnetite is believed to have evolved from the parentdioritic magma and the post magmatic hydrothermal fluids. Conse-quently, it is concluded that the main ore bodies are undoubtedlymagmatic in origin and hydrothermal activity has also played an im-portant role for the late mineralization stage. For example, those al-teration reactions are probably due to a late, hydrothermal phasefrom the same magmatic activity that gave rise to the pegmatitic

Table 1Representative clinopyroxene phenocrysts analyses from core to rim.

Phenocryst-1 Phenocryst-2

Core Rim Core Rim

SiO2 53.05 53.13 52.26 50.53 53.97 52.86 52.22 51.46 50.72 50.24 50.96 49.45 54.25 51.88 49.82 48.52 52.52 52.13 49.06 50.37 49.46TiO2 0.46 0.76 0.5 1.01 0.44 0.59 0.75 0.84 1.25 1.1 1.2 1.15 0.43 0.54 1.13 1.81 0.61 0.68 1.08 1.19 1.54Al2O3 2.89 3.18 3.52 4.93 2.36 3.32 4.02 4.38 4.74 5.26 5.15 4.47 2.04 3.61 5.41 6.41 2.98 5.01 5.05 4.91 4.77Cr2O3 0.16 – 0.14 0.2 0.27 0.68 0.18 0.22 0.19 0.12 0.12 1.42 – – – – – – – – –

FeOT 5.84 6.1 6.92 8.21 5.35 5.64 6.14 6.72 6.85 7.15 8.85 9.17 6.47 5.85 7.61 9.23 6.84 6.89 7.53 7.65 9.78MnO 0.05 0.12 0.13 0.01 0.22 0.12 0.26 0.02 0.14 0.43 0.09 0.3 – 0.3 0.16 0.16 0.06 – – 0.21 0.33MgO 15.36 15.23 14.59 14.25 16.09 15.33 14.58 14.47 14.54 13.52 13.17 12.15 17.25 15.02 13.44 12.49 15.21 14.09 13.71 14.37 12.06CaO 22.11 21.56 20.29 19.34 20.67 21.74 21.66 21.49 21.39 21.91 21.08 20.52 18.43 22.03 21.61 20.93 20.83 21.29 21.58 20.53 20.73Na2O 0.42 0.22 0.5 0.55 0.47 0.37 0.42 0.52 0.42 0.47 0.4 0.68 0.18 0.31 0.3 0.44 0.28 0.54 0.57 0.25 0.35K2O – – 0.03 0.07 – – – – – – – – – – – – 0.04 – 0.04 – –

Total 100.33 100.3 98.88 99.1 99.84 100.66 100.25 100.11 100.25 100.2 101.03 99.32 99.01 99.55 99.48 99.99 99.37 100.62 98.62 99.48 99.01Si4+ 1.941 1.941 1.941 1.884 1.970 1.927 1.915 1.895 1.870 1.861 1.875 1.867 1.987 1.917 1.858 1.815 1.942 1.904 1.850 1.872 1.870Ti4+ 0.013 0.021 0.014 0.028 0.012 0.016 0.021 0.023 0.035 0.031 0.033 0.033 0.012 0.015 0.032 0.051 0.017 0.019 0.031 0.033 0.044Al3+ 0.125 0.137 0.154 0.217 0.102 0.143 0.174 0.190 0.206 0.230 0.223 0.199 0.088 0.157 0.238 0.283 0.130 0.216 0.224 0.215 0.213Cr3+ 0.005 0.000 0.004 0.006 0.008 0.020 0.005 0.006 0.006 0.004 0.003 0.042 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000Fe2+ 0.179 0.186 0.215 0.256 0.163 0.172 0.188 0.207 0.211 0.221 0.272 0.290 0.198 0.181 0.237 0.289 0.211 0.210 0.237 0.238 0.309Mn2+ 0.002 0.004 0.004 0.000 0.007 0.004 0.008 0.001 0.004 0.013 0.003 0.010 0.000 0.009 0.005 0.005 0.002 0.000 0.000 0.007 0.011Mg2+ 0.838 0.829 0.808 0.792 0.876 0.833 0.797 0.794 0.799 0.747 0.722 0.684 0.942 0.827 0.747 0.696 0.838 0.767 0.771 0.796 0.680Ca2+ 0.867 0.844 0.807 0.772 0.808 0.849 0.851 0.848 0.845 0.869 0.831 0.830 0.723 0.872 0.863 0.839 0.825 0.833 0.872 0.817 0.840Na+ 0.030 0.016 0.036 0.040 0.033 0.026 0.030 0.037 0.030 0.034 0.029 0.050 0.013 0.022 0.022 0.032 0.020 0.038 0.042 0.018 0.026K+ 0.000 0.000 0.001 0.003 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.002 0.000 0.002 0.000 0.000Total 3.997 3.978 3.985 3.998 3.979 3.989 3.989 4.002 4.005 4.009 3.992 4.004 3.963 4.001 4.002 4.009 3.987 3.988 4.029 3.996 3.992En 0.44 0.44 0.43 0.43 0.46 0.44 0.43 0.42 0.42 0.40 0.39 0.37 0.50 0.43 0.40 0.37 0.44 0.41 0.40 0.42 0.36Mg# 0.82 0.82 0.79 0.76 0.84 0.83 0.81 0.79 0.79 0.77 0.73 0.70 0.83 0.82 0.76 0.71 0.80 0.78 0.76 0.77 0.69

Phenocryst-3 Phnocryst-4

Core Rim Core Rim

SiO2 53.4 50.11 52.85 51.96 52.82 53.67 50.68 50.7 48.79 50.82 53.2 52.67 50.58 49.9 52.08 52.36 51.81 51.6 51.49 51.66TiO2 0.53 0.1 0.28 0.38 0.6 0.54 0.94 0.96 1.5 1.22 0.54 0.64 0.9 0.93 0.94 0.57 0.83 1.08 0.94 0.65Al2O3 2.79 3 3.72 3.8 3.92 2.11 4.53 5.06 6.25 4.14 2.96 3.7 4.04 5.86 3.58 2.96 3.25 4.72 3.51 2.99Cr2O3 – – – – – – – – – – – – – – – – – – – –

FeO 5.02 5.11 6.6 8.8 5.8 5.29 6.94 7.48 10.3 10.25 5.5 5.66 6.8 8.39 5.83 6.47 6.65 6.73 8.4 9.39MnO 0.17 0.27 0.3 0.16 0.35 0.26 0.33 0.17 0.53 0.38 0.14 0.37 0.03 0.08 0.25 0.18 0.12 0.11 0MgO 15.79 14.12 15.35 13.62 15.13 15.8 13.77 14.07 12.05 12.62 15.14 14.93 14.56 13.11 15 15.42 14.57 14.3 14.24 13.92CaO 21.93 26.02 20.68 20.94 21.38 21.67 21.8 21.52 20.48 20.31 21.54 21.8 21.35 21.01 21.71 20.47 21.08 21.33 20.35 20.41Na2O 0.09 0.27 0.39 0.49 0.36 0.23 0.27 0.2 0.46 0.29 0.54 0.13 0.21 0.14 0.39 0.34 0.29 0.39 0.39 0.6K2O 0.02 0.05 – – 0.08 – – 0.11 – – 0 – – 0.02 – – – – – –

Total 99.56 98.94 100.14 100.28 100.25 99.67 99.18 100.42 100 100.17 99.8 99.67 98.81 99.39 99.61 98.84 98.66 100.27 99.43 99.62Si4+ 1.955 1.887 1.935 1.924 1.929 1.968 1.889 1.871 1.830 1.898 1.952 1.934 1.892 1.862 1.919 1.943 1.932 1.894 1.917 1.930Ti4+ 0.015 0.003 0.008 0.011 0.016 0.015 0.026 0.027 0.042 0.034 0.015 0.018 0.025 0.026 0.026 0.016 0.023 0.030 0.026 0.018Al3+ 0.120 0.133 0.161 0.166 0.169 0.091 0.199 0.220 0.276 0.182 0.128 0.160 0.178 0.258 0.155 0.129 0.143 0.204 0.154 0.132Cr3+ 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000Fe2+ 0.154 0.161 0.202 0.272 0.177 0.162 0.216 0.231 0.323 0.320 0.169 0.174 0.213 0.262 0.180 0.201 0.207 0.207 0.262 0.293Mn2+ 0.000 0.005 0.008 0.009 0.005 0.011 0.008 0.010 0.005 0.017 0.012 0.004 0.012 0.001 0.002 0.008 0.006 0.004 0.003 0.000Mg2+ 0.862 0.793 0.838 0.752 0.824 0.864 0.765 0.774 0.674 0.703 0.828 0.817 0.812 0.729 0.824 0.853 0.810 0.783 0.791 0.775Ca2+ 0.860 1.050 0.811 0.831 0.836 0.851 0.871 0.851 0.823 0.813 0.847 0.857 0.855 0.840 0.857 0.814 0.842 0.839 0.812 0.817Na+ 0.006 0.020 0.028 0.035 0.025 0.016 0.020 0.014 0.033 0.021 0.038 0.009 0.015 0.010 0.028 0.024 0.021 0.028 0.028 0.043K+ 0.001 0.002 0.000 0.000 0.004 0.000 0.000 0.005 0.000 0.000 0.000 0.000 0.000 0.001 0.000 0.000 0.000 0.000 0.000 0.000Total 3.974 4.054 3.991 4.000 3.985 3.979 3.995 4.002 4.007 3.987 3.988 3.973 4.002 3.989 3.991 3.988 3.984 3.988 3.993 4.008En 0.46 0.39 0.44 0.40 0.44 0.45 0.41 0.41 0.36 0.37 0.44 0.44 0.43 0.40 0.44 0.45 0.43 0.42 0.42 0.40Mg# 0.85 0.83 0.81 0.73 0.82 0.84 0.78 0.77 0.68 0.69 0.83 0.82 0.79 0.74 0.82 0.81 0.80 0.79 0.75 0.73

“–” represents below detection limit.

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Table 2Representative clinopyroxenes in the groundmass analyses from core to rim.

Matrix-1 Marix-2 Marix-3 Marix-4

Core Rim Core Rim Core Rim Core Rim

SiO2 51.47 51.73 51.4 51.28 51.42 51.48 51.24 51.45 51.34 51.12 51.67 51.34 51.7 51.3 51.25 51.51TiO2 0.79 0.79 1.04 0.94 0.84 0.92 1.00 0.84 0.82 0.88 0.93 0.97 0.96 0.89 0.98 1.02Al2O3 4.09 3.29 3.36 3.37 4.01 3.62 3.6 3.33 3.47 3.44 3.92 3.6 3.63 3.59 4.23 3.73Cr2O3 0.1 0.07 – 0.06 0.19 0.03 0.01 – 0 0.06 – 0.15 0.04 0.03 0.03 0.18FeO 9.04 8.33 9.13 8.6 8.57 8.45 8.75 8.48 9.34 8.75 9.12 9.1 8.98 9.25 8.46 8.51MnO 0.34 0.35 0.17 0.11 0.15 0.06 0.27 0.23 0.07 0.17 0.17 0.35 – 0.11 0.33 0.31MgO 13.93 14.44 14.16 14.51 14.07 14.26 14.56 14.47 14.26 14.08 14.11 13.85 14.13 14.37 13.96 13.92CaO 20.06 19.75 19.9 20.15 20.92 19.92 20.47 20.36 20.21 20.4 20.22 20.04 20.45 20.36 20.68 20.21Na2O 0.26 0.3 0.31 0.36 0.22 0.33 0.22 0.33 0.27 0.26 0.31 0.33 0.4 0.35 0.4 0.27K2O – – – – – – – – – – – – – – – 0.01Total 100.07 99.04 99.48 99.39 100.39 99.07 100.12 99.48 99.78 99.15 100.44 99.72 100.29 100.24 100.3 99.67Si4+ 1.908 1.930 1.918 1.913 1.901 1.921 1.901 1.917 1.912 1.914 1.909 1.913 1.913 1.903 1.896 1.915Ti4+ 0.022 0.022 0.029 0.026 0.023 0.026 0.028 0.024 0.023 0.025 0.026 0.027 0.027 0.025 0.027 0.029Al3+ 0.179 0.145 0.148 0.148 0.175 0.159 0.157 0.146 0.152 0.152 0.171 0.158 0.158 0.157 0.184 0.163Cr3+ 0.003 0.002 0.000 0.002 0.006 0.001 0.000 0.000 0.000 0.002 0.000 0.004 0.001 0.001 0.001 0.005Fe2+ 0.280 0.260 0.285 0.268 0.265 0.264 0.271 0.264 0.291 0.274 0.282 0.284 0.278 0.287 0.262 0.265Mn2+ 0.011 0.011 0.005 0.003 0.005 0.002 0.008 0.007 0.002 0.005 0.005 0.011 0.000 0.003 0.010 0.010Mg2+ 0.770 0.803 0.788 0.807 0.775 0.793 0.805 0.804 0.792 0.786 0.777 0.769 0.779 0.795 0.770 0.771Ca2+ 0.797 0.790 0.795 0.805 0.828 0.796 0.813 0.813 0.806 0.818 0.800 0.800 0.811 0.809 0.820 0.805Na+ 0.019 0.022 0.022 0.026 0.016 0.024 0.016 0.024 0.019 0.019 0.022 0.024 0.029 0.025 0.029 0.019K+ 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000Total 3.988 3.985 3.990 3.999 3.994 3.985 4.001 3.998 3.998 3.994 3.991 3.990 3.995 4.006 3.999 3.982En 0.41 0.43 0.42 0.42 0.41 0.42 0.42 0.42 0.41 0.41 0.41 0.41 0.41 0.41 0.41 0.41Mg# 0.73 0.76 0.73 0.75 0.75 0.75 0.75 0.75 0.73 0.74 0.73 0.73 0.74 0.73 0.75 0.74

“–” represents below detection limit.

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apatite-bearing iron ores and are mostly of restricted extent and ofsomewhat variable character. In addition, veins with calcite, bariteand pyrite in the ores are also expressions of hydrothermal activity,but these hydrothermal-origin ores are economically unimportantrelative to the magmatic-origin ores.

In summary, there is little doubt that the massive ores and brecci-ated ores in the Gushan magnetite–apatite deposit are magmatic inorigin and crystallized from iron oxide melts, and the subeconomicores such as stockwork and skeleton ores were formed by hydrother-mal fluids (Chang et al., 1991; Hou et al., 2009; Mao et al., 2006; Songet al., 1981; Yu and Mao, 2002). Similar magmatic features of ironores of Kiruna-type have also been recognized in the El Laco volcanoin Chile (Park, 1961), and their ancient equivalents at Kiruna in Swe-den (Frietsch and Perdahl, 1995; Martinsson, 1997; Martinsson andWeihed, 1999).

Table 3Representative plagioclase phenocrysts analyses from core to rim.

Phenocryst-1 Phenocryst-2 Phenocryst-3

Core Innerrim

Rim Core Innerrim

Rim Core Innerrim

Rim

SiO2 49.38 50.27 50.78 51.08 50.5 51.22 49.57 50.4 53.72Al2O3 31.69 30.63 29.88 31.19 29.97 29.35 31.24 31.16 29.25FeO 0.46 1.02 0.89 1.07 0.94 1.41 0.96 0.55 0.23CaO 13.67 13.09 12.67 13.18 12.77 12.36 13.59 12.98 10.57Na2O 2.77 3.73 3.99 3.37 3.66 3.57 3.03 3.48 4.74K2O 0.49 0.31 0.49 0.49 0.44 0.38 0.3 0.16 0.6Total 98.69 99.49 98.97 100.38 98.43 98.51 98.85 98.95 99.27Si4+ 9.144 9.278 9.399 9.293 9.380 9.505 9.177 9.284 9.787Al3+ 6.916 6.663 6.518 6.687 6.561 6.418 6.816 6.764 6.280Fe2+ 0.071 0.157 0.138 0.163 0.146 0.219 0.149 0.085 0.035Ca2+ 2.712 2.588 2.513 2.569 2.541 2.457 2.695 2.561 2.063Na+ 0.994 1.335 1.432 1.189 1.318 1.284 1.087 1.243 1.674K+ 0.116 0.073 0.116 0.114 0.104 0.090 0.071 0.038 0.139Total 19.953 20.094 20.115 20.014 20.050 19.973 19.995 19.974 19.979An 0.71 0.65 0.62 0.66 0.64 0.64 0.70 0.67 0.53Ab 0.26 0.33 0.35 0.31 0.33 0.34 0.28 0.32 0.43Or 0.03 0.02 0.03 0.03 0.03 0.02 0.02 0.01 0.04

“–” represents below detection limit.

6.3. Petrographic evidence for liquid immiscibility

A model of fractional crystallization has been proposed to inter-pret the generation of the other magmatic iron oxide deposits, e.g.,the Panzhihua V–Ti–Fe deposit in Emeishan large igneous province(e.g., Zhang et al., 2009). Abundant Fe–Ti oxide inclusions in cumulusolivine from the Panzhihua intrusions show evidence for early crys-tallization of Fe–Ti oxides in ferrobasaltic magma, which is consistentwith the interpretation that the stratiform oxide ores in the Panzhi-hua intrusions formed by accumulation of Fe–Ti oxide crystalswhich are formed by fractional crystallization rather than immiscibleoxide melt. Unlike the Gushan deposit in the Ningwu basin, those inthe Emeishan large igneous province occur as iron beds associatedwith the gabbros, and are generally concentrated in the lower partsof the intrusion (e.g., Zhou et al., 2005) while the ore bodies in the

Phenocryst-4 Phenocryst-5 Phenocryst-6

Core Innerrim

Rim Core Innerrim

Rim Core Innerrim

Rim

50.86 51.5 52.72 49.68 50.94 49.84 50.99 50.56 5232.21 31.07 31.69 31.63 31.49 31.46 31.33 31.22 29.820.77 0.77 0.83 0.8 0.91 1.05 0.71 0.56 1.0313.34 12.96 9.94 14.18 13.37 13.42 13.29 13.44 12.282.93 3.64 3.25 2.86 3.53 3.18 3.51 3.44 3.820.33 0.31 1.18 0.28 0.37 0.32 0.37 0.41 0.5100.97 100.65 99.97 99.55 100.73 99.37 100.4 99.86 99.839.213 9.354 9.547 9.132 9.247 9.175 9.280 9.257 9.5226.877 6.651 6.763 6.852 6.736 6.825 6.720 6.737 6.4350.117 0.117 0.126 0.123 0.138 0.162 0.108 0.086 0.1582.589 2.522 1.928 2.792 2.600 2.647 2.591 2.636 2.4091.029 1.282 1.141 1.019 1.242 1.135 1.238 1.221 1.3560.076 0.072 0.273 0.066 0.086 0.075 0.086 0.096 0.11719.901 19.997 19.778 19.984 20.049 20.018 20.023 20.033 19.9970.70 0.65 0.58 0.72 0.66 0.69 0.66 0.67 0.620.28 0.33 0.34 0.26 0.32 0.29 0.32 0.31 0.350.02 0.02 0.08 0.02 0.02 0.02 0.02 0.02 0.03

Table 4Representative alkali feldspar in the groundmass analyses from core to rim.

Matrix-1 Matrix-2 Matrix-3 Matrix-4 Matrix-5

Core Rim Core Rim Core Rim Core Rim Core Rim

SiO2 64.54 64.66 64.17 63.99 65.06 65.64 65.63 64.67 64.45 64.15Al2O3 19.87 19.29 18.75 20.49 19.42 18.98 18.7 19.11 19.53 19.08FeO 0.38 – 0.19 – 0.26 0.22 0.19 0.11 – 0.22CaO 0.78 0.45 0.62 1.28 0.7 0.53 0.44 0.54 0.73 0.76Na2O 5.91 4.14 3.85 5.63 5.62 5.34 5.76 4.27 4.39 4.11K2O 8.4 10.45 10.79 8.84 8.83 9.46 8.82 10.67 10.56 10.94Total 100.03 99.33 98.9 100.36 99.89 100.24 99.63 99.43 99.93 99.47Si4+ 11.705 11.852 11.875 11.591 11.796 11.881 11.922 11.843 11.772 11.798Al3+ 4.247 4.167 4.089 4.374 4.149 4.049 4.003 4.124 4.204 4.136Fe2+ 0.058 0.000 0.029 0.000 0.039 0.033 0.029 0.017 0.000 0.034Ca2+ 0.152 0.088 0.123 0.248 0.136 0.103 0.086 0.106 0.143 0.150Na+ 2.078 1.471 1.381 1.977 1.975 1.874 2.029 1.516 1.555 1.465K+ 1.943 2.443 2.547 2.042 2.042 2.184 2.044 2.493 2.460 2.567Total 20.182 20.022 20.045 20.232 20.138 20.124 20.112 20.099 20.134 20.150An 0.04 0.02 0.03 0.06 0.03 0.02 0.02 0.03 0.03 0.04Ab 0.50 0.37 0.34 0.46 0.48 0.45 0.49 0.37 0.37 0.35Or 0.47 0.61 0.63 0.48 0.49 0.52 0.49 0.61 0.59 0.61

“–” represents below detection limit.

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Gushan deposit exclusively occur at the top of the intrusion due to thelow density of the ore-bearing magma with enormous volatile con-tents, as indicated by many miarolitic cavities in the massive oresand brecciated ores. The typical flow banding structure in the massiveores is good evidence for magmatic flow. In addition, the absence ofmagnetite and hematite associated with silicates and accumulation-texture precludes the possibility of fractional crystallization. Alter-nately, all features indicate that the massive ores likely formed froma cooling immiscible magma phase.

The role of liquid immiscibility in the formation of igneous rocks is awidely discussed subject (Thompson et al., 2007). It is proposed thatsuch processes are involved in the genesis of carbonatites (Koster vanGroos andWyllie, 1963 and many later contributions) and of Fe-rich re-sidual melts in terrestrial and lunar basalts (Roedder, 1984 and others),especially the Skaergaard intrusion in East Greenland (Veksler, 2009and references therein), the nelsonite from the Antauta subvolcanic cen-ter, Peru (Clark and Kontak, 2004, Kolker, 1982) and the typical Kiruna-type iron oxide deposits (e.g., Lyons, 1988). However, considering thatthe signs of liquid immiscibility are notoriously elusive (Bowen, 1928;Roedder, 1979), especially since the iron ores experienced intensive hy-drothermal activity, and no macroscopic evidence for immiscibility hasbeen found, such as ocelli, globular structures (Markl, 2001), spheroidal

Fig. 4. Plot of V2O5 vs. TiO2 content (wt.%) in magnetite formed by different genesis.Compositional fields for El Laco and Kiruna are from Nyström and Henriquez (1994);other data are from Ren (1991).

bodies (Ferguson, 1964; Sørensen et al., 2003) or ovoid bodies (Bussenand Sakharov, 1971) in the rocks. Therefore, the directly powerful evi-dence of liquid immiscibility and following magmatic injection of ironoxidemagma is still absent. Accordingly geochemical indicators are like-ly to bemore useful. One of the geochemical fingerprints of themagmat-ic liquid immiscibility is the strong enrichment in P, Ti, REE and HFSE,and higher CaO/Al2O3 ratios in the massive ores (Song et al., 1981;Zhao, 1993a,b) which are the geochemical hallmarks of Fe-rich immisci-ble liquids (Veksler et al., 2006). However, it should be pointed out thatany remnants of the silicate-rich liquid in the iron-rich counterpart willobscure the geochemical indicators and the origin of the rock by immis-cibility. Thus we need to consider whether liquid immiscibility occurredby an alternative approach.

Investigation of mineral textures and zoning as evidence for opensystem processes during magmatic evolution have always been a cen-terpiece of petrological studies and have provided some of the best ev-idence for magma mixing (e.g., Andrews et al., 2008), magma recharge(e.g., Turner et al., 2008), and crustal contamination (e.g., Anderson,1976) for many decades. Asminerals record textural and compositionalchange of the magmatic system, they preserve in their crystal growthzoning awealth of information regarding their past history ofmagmaticprocesses and compositions (Ginibre et al., 2007; Streck, 2008). Thus, aseparate record of liquid immiscibility should also be recorded in com-positionally zoned phenocrysts because major and trace element com-positions of the crystallizing phenocrysts are strong functions of melttemperature and composition (e.g., Bachman and Dungan, 2002; Corteset al., 2005; Gardner et al., 1995; Holland and Blundy, 1994; Holtz et al.,2005; Tsuchiyama and Takahashi, 1983). The study byHou et al. (2010)suggests that clinopyroxene and plagioclase were the major fractionat-ing phases during the magmatic evolution of the Ningwu dioritic por-phyries. As a result, Mg will behave compatibly and decrease in themore evolved magma. Furthermore, other elements, such as Fe, willnot be incorporated preferentially in the clinopyroxene and their con-centrationswill increase in the evolving liquid. Additionally, theAb con-tents will progressively increase and the An contents should decreaseresponsively. If Fe-rich liquid immiscibility occurred during themagma evolution, the Fe contents in the clinopyroxene phenocrystswill dramatically decrease whereas the compositions of plagioclasephenocrystswill display a normal evolution trend. Interestingly, the ob-served chemical variations in the interiors of the analyzed clinopyrox-ene phenocrysts are consistent with such an evolutionary trend. Thesharply decreased concentration of Fe and elevated Mg in the middleparts of the clinopyroxene phenocrysts reflect changes that likely oc-curred in the evolving liquid (Fig. 5). This sharp reverse zoning can be

Fig. 5. Compositional transects of selected clinopyroxene phenocrysts from Gushan dioritic porphyry. Rim to core compositional profiles of four phenocrysts were determined byEMP analyses. The start and end of clinopyroxene transects were located using Back-Scattered Electron (BSE) images collected with the microprobe.

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explained by two causes: 1) a mafic magma recharge in the evolvingmagma chamber beneath the dioritic porphyries (Ginibre and Wörner,2007; Kamenov et al., 2005); 2) the Fe-rich melt separated from theevolvingmagma by liquid immiscibility, which is a logical consequence.If the mafic magma recharge episode is chosen to be responsible for theexotic reverse zoning, it requires the witness plagioclase phenocrystsand/or the feldspar in the groundmass also to respond to it composi-tionally, i.e., elevated An contents (Streck, 2008; Tepley et al., 2000).However, the major element composition of plagioclase phenocrystsfrom core to rim shows that the An content generally decreases gradu-ally. Given the more alkaline affinity of the feldspar in the groundmass,we can preclude the possibility that the reverse zoning in clinopyroxenephenocrysts could be caused by mafic magma recharge. Thus, the mostreasonable explanation for the origin of these reverse zoning patterns inclinopyroxene phenocrysts is that an immiscible Fe-rich liquid separat-ed from the extreme iron-rich magmas during differentiation. Thesemagmas were derived from a common parent basaltic melt that experi-enced fractionation of plagioclase and clinopyroxene during ascent tothe surface, which led to the extreme enrichment of iron in a highlyevolved magma (e.g., Hou et al., 2010).

6.4. Triggering of liquid immiscibility

Our former investigation on the Gushan dioritic porphyry (Hou etal., 2010) shows that it is derived from a common parent basalticmelt, which was produced by partial melting of the enriched litho-spheric mantle, and the parental melt experienced assimilation ofthe Yangtze craton upper crust, accompanying the fractionation ofplagioclase and clinopyroxene as a Fenner trend during ascent tothe surface, which led to the extreme enrichment of iron in a highlyevolved magma. The fractionation of the gabbroic assemblage is evi-denced by the presence of clinopyroxene and labradorite phenocrysts(An=~70). Besides, the presence of labradorite phenocrysts(An=~70) suggests that they were crystallized from basalticmagmas because the dioritic rocks generally contain andesine ratherthan labradorite (Garavaglia et al., 2002). Moreover, some gabbroicrocks are present at the bottom of the dioritic stocks in Gushan(Tang et al., 1998; Xu and Xing, 1994).

Many previous studies demonstrated experimentally that addi-tions of even minor amounts of phosphorous appear to have signifi-cant effects on liquid immiscibility, also in anhydrous systems (e.g.,Bogaerts and Schmidt, 2006; Philpotts, 1967; Ryerson and Hess,1980; Suk 1998; Visser and Koster van Groos, 1979a,b,c). Sillitoeand Burrows (2002) proposed that the El Laco magnetite lava flowin northern Chile may owe its formation to phosphorous-enhancedimmiscibility and fractionation effects. Therefore, the convergenceof experimental and field evidence suggests that iron oxide oremagma can indeed form in nature, and that the textural evidencefor their existence should be taken at face value. Consequently, inthese deposits, we prefer that the addition of minor phosphorousinto the evolving ore-bearing magma is most likely the trigger for liq-uid immiscibility in spite of the fact that that some other conditionsare also capable of facilitating development of an immiscible oxidemelt, e.g., the involvement of CO2 (Zhou et al., 2005). With regardto the source for the phosphorus, Yu et al. (2008) argued that forma-tion of the apatites from magnetite–apatite tabular bodies with inter-growth texture was due to crustal contamination mainly based on thenarrow variation range of 87Sr/86Sr values of apatites from 0.706326to 0.707577, similar to those of the volcanic and subvolcanic rocksand higher than that of a typical mantle-derived magmatic apatite(i.e., about 0.7040). This is also revealed by the Pb isotope composi-tions of apatite, magnetite and pyrite (e.g., Ma et al., 2006), the Sr–Nd isotope compositions of the dioritic porphyries (e.g., Hou et al.,2010 and reference therein) and the Hf isotope compositions of thezircon in the dioritic porphyries (Hu and Jiang, 2010). However,there are still two possible sources for the phosphorus, i.e., the basinal

brine and the strata. Since the P content is extremely low in basinalbrine which chiefly contains Na+, K+, Ca2+, Mg2+, Cl− and sulfate(Hu and Hu, 1991), we can conclude that the minor addition of P ispredominantly introduced by crustal contamination from P-rich stra-ta. There are many P2O5-rich beds visible in the LYRV from Cambrianto Permian strata along the Yangtze River, which contain somephosphorus-rich layers such as the marlstone rocks of lower-middleJurassic Xiangshan Group with average P2O5 concentration up to6.93 wt.% (Zhao, 1993a,b). Thus, it is possible that the Fe-rich dioriticmagma which formed by fractionation of Fe-minerals, i.e., clinopyr-oxene and plagioclase, had been contaminated by some phosphorusmaterials during magma ascent, as suggested by the geochemicalcharacteristics of the dioritic stocks. As a result, the ore-bearing dio-ritic magma split into Fe-poor silicate melt and iron oxide melt by liq-uid immiscibility due to addition of minor phosphorus from thecountry rock.

6.5. Post-magmatic hydrothermal activities

The post-magmatic hydrothermal alteration in the Gushan irondeposit mainly occurs along the mineralized fractured zones. Howev-er, as a whole, the alteration is relatively weak. The main alterationtypes include kaolinization, silicification and carbonatization. The ex-tent of alteration depends on the composition of the original rocksand the distance from the mineralized zone. Weak silicification andkaolinization occurred in the dolomite limestone of the XiangshanGroup and the mud shale and siltstone of Huangmaqing Fm., but rep-resentative skarn minerals have not been recognized. Due to the ther-modynamic effect of the magma, metamorphism occurred in thecountry rocks, which led to recrystallization of limestone and forma-tion of quartzite from quartz sandstone. The skeleton ore consisting ofrandomly-oriented coarse martitized magnetite crystals in the micro-crystalline groundmass of martite and hematite with occasionalquartz and/or chalcedony infill, and the stockwork ores have beeninterpreted to be of hydrothermal origin (Gu and Ruan, 1988).

The dioritic porphyry has undertaken intensive hydrothermal al-teration, which comprises a light-color alteration zone and a darkalteration zone. The former includes silicification and kaolinization,which occurs at near ore bodies. The silicification led to elevatedSiO2 contents of dioritic porphyry, up to 68.32 wt.% (Zhai et al.,1992) and subordinary quartz in the porphyry. The kaolinization ischaracterized by the plagioclases and feldspars replaced by kaoliniteand dramatically varied chemical compositions. The dark-coloralteration zone is characterized by the replacement of chlorite, andcarbonate minerals. Of those alteration types, silicification andkaolinization is the most important, because they are the indicatorof mineralization.

6.6. A proposed metallogenetic model

As discussed above, the Fe-rich melts separated from the ore-bearing magma by liquid immiscibility and were then injected alongfractures and spaces between the dioritic intrusions and wall-rocks.However, it has been pointed out in some previous works that thedense Fe-rich liquids are not likely to rise to the top of a magmachamber or erupt (e.g., Veksler et al., 2006). Therefore, they wouldstay behind in magmatic chambers, or migrate downwards in mag-matic plumbing systems. This is the main confusion, in our view,that workers who favored the magmatic in origin of exotic Kiruna-type mineralization may encounter, and partly the reasons why twoconflicting theoretical hypothesis, i.e., magmatic and hydrothermal,were proposed (e.g., Institute of Geochemistry of Chinese Academyof Sciences, 1987; Lin et al., 1983; Lu et al., 1990; Ningwu ResearchGroup, 1978).

First of all, the enrichment of volatiles in the ore-bearing dioriticmagmas has been reached by fractionation of anhydrous minerals

343T. Hou et al. / Ore Geology Reviews 43 (2011) 333–346

(e.g., clinopyroxene and plagioclase), as indicated by the fact that thereare stomas, tubes, and miarolitic and amygdaloidal structures in themassive ores. During the magma ascent, the magmas were contaminat-ed by the phosphor layers and nodules and the carbonate layers (CO2)elevated the concentrations of volatiles. When the ore-forming magmawas emplaced near the surface, thefluxing agents, mainly the phosphor,facilitated development of the immiscible oxide melt. The intensive oc-currence of fluorapatites in massive ore bodies suggests that the volatilephases are selectively partitioned into these immiscible Fe-rich meltsand reach relatively higher concentrations (e.g., Kolker, 1982) duringthe liquid immiscibility. This is evidenced by the experiment that P2O5

is strongly enriched in the Fe-rich melt (DPM/F between 3.8 and 8.9 in

the system Fe2SiO4–KAlSi3O8–SiO2–P2O5; Bogaerts and Schmidt, 2006)and is able to lower the liquidus temperatures of Fe-rich melts (Clarkand Kontak, 2004). The fact that crystallization of magnetite is late inthe emplacement of dioritic porphyries could facilitate a relatively lon-ger period of contamination by the reaction between Fe-rich melts andthe country rocks such as carbonate and gypsum.Moreover, the concen-trations of other volatiles, e.g., F and Cl, could also have been elevated inthe Fe-richmelts because their partition coefficients in apatite are higherthan those in silicate melts (Webster et al., 2009). Therefore, at the

Fig. 6. Proposed metallogenic model of the Kiruna-type iron oxide

magmatic temperatures envisioned, an immiscible Fe-rich melt withenormous volatile contents is believed to have evolved from the parentore-bearing dioritic magma by the introduction of phosphorous and theCO2 from carbonatewall rocks. This Fe- and volatile-richmelt rose to thetop of the magma chamber forcefully due to its low density.

Additional evidence for rapid cooling, and an indication of thehigh crustal level emplacement of the dioritic stocks in the area, isthe rapid transition from ductile to brittle deformation. Furthermore,taking the hypabyssal environment into consideration, fluid exsolu-tion developed by decompression during ore-forming magma ascent,and the growth of bubbles constitutes the driving force behind risingmagmas. Boiling occurred due to the suddenly decreasing pressure,and the internal fluid pressures became suddenly larger than thelithostatic load. Eventually, new bubble formation led to an explosionnear the surface, resulting in the immediate fragmentation of the roofof the intrusion and wall-rocks.

These fragments were soon cemented by Fe oxide melts, formingbrecciated ores. The residual oxide melts filled in the small faultsand fissures, leading to the formation of massive ores (Fig. 6) anddendritic magnetite. The post-magmatic hydrothermal activitiescaused extensive contact metasomatism in the vicinity of contact

deposits in Ningwu basin. See text for additional explanation.

344 T. Hou et al. / Ore Geology Reviews 43 (2011) 333–346

zones between stocks and wall-rocks and along the fissures, whichare possibly related to later tectonic activities. The grade enrichmentand superimposed reformation of the pre-existing magmatic ores arerelated to these processes. The hydrothermal activities can also ac-count for considerable amounts of stockwork and skeleton ores. Thehydrothermal fluids were possibly formed by the mixture of magmat-ic fluids, meteoric water and the basinal brine, which is probably ca-pable of carrying significant volumes of iron in the form of Na –Fe –

Cl complexes. The fluids are believed to give rise to the pegmatiticapatite-bearing iron ores, the transition of partial magnetite tohematite.

Therefore, based on these investigations, we propose the followingnew model (Fig. 6):

1) Basaltic magma differentiated under a low oxygen fugacity envi-ronment as a Fenner trend, which led to the extreme enrichmentof iron in a highly evolved magma (ferrodioritic composition);

2) Phosphorus was added into the magma by contamination fromthe upper crust;

3) Volatile-rich ore magma formed by liquid immiscibility, probablycaused by the addition of phosphorous into the ore-bearingmagma, consisting of ferrodioritic melts. The volatile-rich oremagma floats up into the roof of the intrusion;

4) Ore magma injected into the contact zone between dioritic stocksand wall-rocks, forming massive, brecciated, and some bandedores due to volatile exsolution, and minor disseminated ores dueto crystallization of magnetite in the relatively Si-rich melts;

5) The extensive post-magmatic hydrothermal activities caused con-tact metasomatism and formed stockwork and skeleton ores andlarge scale alteration.

In summary, this new model does not need a direct link with theiron liquid at the base of the magma chamber and a channel or feederbelow the breccia ore. The phosphorous sourced from both magmaand strata preferentially partitioned into the iron ore magma whichhas been experimentally proven and also has been evidenced by thehigh content of P in the massive ores. Finally, the widespread alter-ation is typically of hydrothermal origin. Such a magmatic injectionorigin for the Gushan deposit in the Ningwu basin is compatiblewith present knowledge of the deposit. In addition, it has room forpartial acceptance of hydrothermal replacement, for replacement tex-ture is common because of the volatile-rich nature of the magma.

7. Concluding remarks

Based on field andmineral chemistry evidence, combined with thepetrological and geochemical features, a subvolcanic origin meritsconsideration for the Gushan deposit. However, it should be remem-bered that replacement textures are all integral parts of this model.The reverse zoning patterns in clinopyroxene phenocrysts is in accor-dance with an immiscible Fe-rich liquid separated from the extremeiron-rich magmas during differentiation under the circumstancesthat the mafic magma recharge had been precluded. The liquid im-miscibility which was triggered by the addition of phosphorus bycrustal contamination can form the uniquemineralization, i.e., the ex-istence of stomas, tubes, and miarolitic and amygdaloidal structureswhich are filled by quartz and calcite crystals in massive ores, thepresence of “ore breccia” and the flow banding structure in massiveore. The occurrence of massive iron ore bodies can be satisfactorilyexplained by the immiscible Fe-rich melt with enormous volatile con-tents that were driven to the top of chamber due to their low density.The hot and volatile-rich iron ore magma all injected along fracturesand spaces between the dioritic intrusions and wall-rocks, and ledto an explosion near the surface, resulting in the immediate fragmen-tation of the roof of the intrusion and wall-rocks, forming brecciatedores. Generally, other types of ores can be considered to result frompost-magmatic hydrothermal activities.

Acknowledgments

The authors are grateful to the management of the Ma'anshan Ironand Steel Group and Shanghai Baoshan Iron & Steel Group for logisticalsupport during fieldwork at the Ningwu area. Financial support for thiswork was supported by “The Fundamental Research Funds for the Cen-tral Universities”, the Projects (2012CB416805 and 2007CB411405) ofthe State Key Fundamental Program, Key Project of Chinese Ministryof Education (No. B07039), the National Natural Science Foundation ofChina (Nos. 40925006 and 40821061), the 111 Project (B07011), andPCSIRT.

References

Aftabi, A., Mohseni, S., Babeki, A., Azaraein, H., 2009. Fluid inclusion and stable isotopestudy of the Esfordi apatite–magnetite deposit, central Iran—a discussion. Econ.Geol. 104, 137–140.

Anderson, A.T., 1976. Magma mixing: petrological process and volcanological tool. J.Volcanol. Geotherm. Res. 1, 3–33.

Andrews, B.J., Gardner, J.E., Housh, T.B., 2008. Repeated recharge, assimilation, and hybridi-zation in magmas erupted from El Chichón as recorded by plagioclase and amphibolephenocrysts. J. Volcanol. Geotherm. Res. 175, 415–426.

Bachman, O., Dungan, M.A., 2002. Temperature-induced Al-zoning in hornblendes ofthe Fish Canyon magma, Colorado. Am. Mineral. 87, 1062–1076.

Barton, M.D., Johnson, D.A., 1996. Evaporitic-source model for igneous-related Feoxide–(REE–Cu–Au–U) mineralization. Geology 24, 259–262.

Barton, M.D., Johnson, D.A., 2000. Alternative brine sources for Fe oxide(–Cu–Au) systems:implications for hydrothermal alteration andmetals. In: Porter, T.M. (Ed.), Hydrothermaliron oxide copper–gold and related deposits: a global perspective. Australian MineralFoundation, Adelaide, pp. 43–60.

Bogaerts, M., Schmidt, M.W., 2006. Experiments on silicate melt immiscibility in thesystem Fe2SiO4–KAlSi3O8–SiO2–CaO–MgO–TiO2–P2O5 and implications for naturalmagmas. Contrib. Mineral. Petrol. 152, 257–274.

Bookstrom, A.A., 1977. The magnetite deposits of El Romeral, Chile. Econ. Geol. 72,1101–1130.

Bookstrom, A.A., 1995. Magmatic features of iron ores of the Kiruna Type in Chile andSweden: ore textures and magnetite geochemistry—a discussion. Econ. Geol. 90,469–473.

Bowen,N.L., 1928. The Evolution of the Igneous Rocks. PrincetonUniversity Press, Princeton,New Jersey, p. 334.

Bussen, I.V., Sakharov, A.S., 1971. Luyavurt ovoidophyre, a liquation product of alkalicmagma. Dokl. Acad. Sci. U.S.S.R Earth Sci. Sect. 197, 170–173.

Chang, Y.F., Liu, X.P., Wu, C.Y., 1991. The Copper–Iron Belt of the Lower and MiddleReaches of the Changjiang River. Geological Publishing House, Beijing. 363 pp.(in Chinese with English abstract).

Clark, A.H., Kontak, D.J., 2004. Fe–Ti–P oxide melts generated through magma mixingin the Antauta Subvolcanic Center, Peru: implications for the origin of Nelsoniteand iron oxide-dominated hydrothermal deposits. Econ. Geol. 99, 377–395.

Cortes, J.A., Wilson, M., Condliffe, E., Francalanci, L., Chertkoff, D.G., 2005. The evolution ofthemagmatic systemof Stromboli Volcano during the Vancori Period (26–13.8 k.y.). J.Volcanol. Geotherm. Res. 147, 1–38.

Fan, Y., Zhou, T.F., Yuan, F., Zhang, L.J., Qian, B., Ma, L., David, R.C., 2010. Geochronologyof the diorite porphyrites in Ningwu basin and their metallogenic siginificances.Acta Pet. Sin. 26 (9), 2715–2728 (in Chinese with English abstract).

Ferguson, J., 1964. Geology of the Ilímaussaq alkaline intrusion, South Greenland. Bulletin-Grønlands Geologiske Undersøgelse, 39 (82 pp., also Meddelelser om Grønland 172(4)).

Förster, H., Jafarzadeh, A., 1994. The Bafqmining district in central Iran a highlymineralizedinfracambrian volcanic field. Econ. Geol. 89, 1697–1721.

Frietsch, R., 1978. On the magmatic origin of iron ores of the Kiruna type. Econ. Geol.73, 478–485.

Frietsch, R., 1991. New ore types in the northern part of the Fennoscandian shield. Ge-ology 113, 46–48.

Frietsch, R., Perdahl, J.A., 1995. Rare earth elements in apatite and magnetite in Kiruna-type iron ores and some other iron ore types. Ore Geol. Rev. 9, 489–510.

Gao, D.M., Zhao, Y.J., 2008. Rerecognition of porphyrite iron ore deposit. Geol. Anhui 18,164–168 (in Chinese with English abstract).

Garavaglia, L., Bitencourt, M.F., Nardi, L.V.S., 2002. Cumulatic diorites related topostcollisional, Brasilianol Pan-African mafic magmatism in the Vila Nova Belt,Southern Brazil. Gondwana Res. 5, 519–534.

Gardner, J.E., Rutherford, M., Carey, S., Sigurdsson, H., 1995. Experimental constraintson pre-eruptive water contents and changing magma storage prior to explosiveeruptions of Mount St. Helens Volcano. Bull. Volcanol. 57, 1–17.

Geological Survey of Industrial Ministry, 1935. Records of Iron Deposits in the LowerReaches of Yangtze River. (in Chinese).

Gibson, S.A., 2002. Major element heterogeneity in Archean to recent mantle plumestarting-heads. Earth Planet. Sci. Lett. 95, 59–74.

Ginibre, C.,Wörner, G., 2007. Variable parentmagmas and recharge regimes of the Parinacotamagma system (N. Chile) revealed by Fe, Mg and Sr zoning in plagioclase. Lithos 98,118–140.

Ginibre, C., Wörner, G., Kronz, A., 2007. Crystal zoning as an archive for magma evolution.Elements 3, 261–266.

345T. Hou et al. / Ore Geology Reviews 43 (2011) 333–346

Gleason, J.D., Marikos, M.A., Barton, M.D., Johnson, D.A., 2000. Neodymium isotopic study ofrare earth element sources andmobility in hydrothermal Fe oxide (Fe–P–REE) systems.Geochim. Cosmochim. Acta 64, 1059–1068.

Gu, L.X., Ruan, H.C., 1988. On the hydrothermal origin of the Gushan iron ore deposit inAnhui Province. Miner. Deposits 7 (2), 21–30 (in Chinese with English abstract).

Hildebrand, R.S., 1986. Kiruna-type deposits: their origin and relationship to intermediatesubvolcanic plutons in the Great Bear Magmatic Zone, northwest Canada. Econ. Geol.81, 640–659.

Hitzman, M.W., 2000. Iron oxide–Cu–Au deposits: what, where, when, and why. In:Porter, T.M. (Ed.), Hydrothermal Iron Oxide Copper–gold & Related Deposits: AGlobal Perspective, vol. 2. PGC Publishing, Adelaide, Australia, pp. 9–25.

Hitzman, M.W., Oreskes, N., Einaudi, M.T., 1992. Geological characteristics and tectonicsetting of Proterozoic iron oxide (Cu–U–Au–REE) deposits. Precambrian Res. 58,241–287.

Holland, T., Blundy, J., 1994. Non-ideal interactions in calcite amphiboles and theirbearing on amphibole-plagioclase thermometry. Contrib. Mineral. Petrol. 116,433–447.

Holtz, F., Sato, H., Lewis, J., Behrens, H., Nakada, S., 2005. Experimental petrology of the1991–1995 Unzen dacite, Japan: part I, phase relations, phase compositions andpre-eruptive conditions. J. Petrol. 46, 319–337.

Hou, K.J., Yuan, S.D., 2010. Zircon U–Pb age and Hf isotopic composition of the volcanicand sub-volcanic rocks in the Ningwu basin and their geological implications. ActaPetrol. Sin. 26, 888–902 (in Chinese with English abstract).

Hou, T., Zhang, Z.C., Du, Y.S., Li, S.T., 2009. Geology of theGushan iron oxide deposit associatedwith dioritic porphyries, eastern Yangtze craton, SE China. Int. Geol. Rev. 51, 520–541.

Hou, T., Zhang, Z.C., Encarnacion, J., Du, Y.S., Zhao, Z.D., Liu, J.L., 2010. Geochemistry ofLate Mesozoic dioritic porphyries associated with Kiruna-style and strataboundcarbonate-hosted Zhonggu iron ores, Middle-Lower Yangtze Valley, EasternChina: Constraints on petrogenesis and iron sources. Lithos 119, 330–344.

Hu, W.X., Hu, S.X., 1991. Na source of albitization in Ningwu-type iron-sulfur depositsof Ningwu-Luzong area. Contrib. Geol. Miner. Resour. Res. 6 (2), 36–46 (in Chinesewith English abstract).

Hu, J.P., Jiang, S.Y., 2010. Zircon U–Pb dating and Hf isotopic compositions of porphyritesfrom the Ningwu basin and their geological implications. Geol. J. China Univ. 16,294–308 (in Chinese with English abstract).

Institute of Geochemistry of Chinese Academy of Sciences, 1987. Ore-forming Mechanismof Ningwu Type Iron Deposits. Science Press, Beijing, pp. 62–76 (in Chinese).

Jami, M., Dunlop, A.C., Cohen, D.R., 2007. Fluid inclusion and stable isotope study of theEsfordi apatite-magnetite deposit, central Iran. Econ. Geol. 102, 1111–1128.

Jami, M., Dunlop, A.C., Cohen, D.R., 2009. Fluid inclusion and stable isotope study of theEsfordi apatite–magnetite deposit, central Iran—a reply. Econ. Geol. 104, 140–143.

Jiang, S.Y., Ma, F., Zhao, K.D., Jiang, Y.H., Ling, H.F., Ni, P., 2006. Origin of iron deposits inthe Ningwu volcanic basin, Lower Yangtze River district, China: geochemical andisotopic evidence. Goldschmidt Conference Abstracts A293.

Kamenov, G.D., Perfit, M.R., Jonasson, I.R., Mueller, P.A., 2005. High-precision Pb isotopemeasurements reveal magma recharge as a mechanism for ore deposit formation:examples from Lihir Island and Conical seamount, Papua New Guinea. Chem. Geol.219, 131–148.

Kolker, A., 1982. Mineralogy and geochemistry of Fe-Ti oxide and apatite (nelsonite)deposits and evaluation of the liquid immiscibility hypothesis. Econ. Geol. 77,1146–1158.

Koster van Groos, A.F., Wyllie, P.J., 1963. Experimental data bearing on role of liquidimmiscibility in genesis of carbonatites. Nature 199, 801–802.

Li, B.L., Xie, Y.H., 1984. The origin, classification, and genesis model of the Ningwu Typeiron deposits in the Nanjing and Wuhu area. Sci. Sin. 1, 80–86 (in Chinese).

Lin, C.X., Zhang, Z.R., Bai, Z.H., 1983. On the conditions of formation of Gushan iron deposit.Geochimica 1, 15–24 (in Chinese with English abstract).

Lin, G., Zhu, C.L., Xu, D.R., 2010. Discussion on metallogenetic model of the southernNingwu porphyrite iron deposit and deep prospecting method. Geotectonica etMetallogenia 34, 368–377 (in Chinese with English abstract).

Lu, B., Hu, S.X., Lin, Y.S., Ye, S.Q., 1990. A study on origin and ore-forming model ofNingwu type iron deposits. Mine. Deposits 9, 13–24 (in Chinese with Englishabstract).

Lü, B.Q., Wang, H.G., Hu, W.S., Shen, W.F., Zhang, Y.L., 2004. Relationship between Paleozoicupwelling facies and hydrocarbon in southeastern marginal Yangtze Block. Mar. Geol.Quat. Geol. 24 (4), 29–35 (in Chinese with English abstract).

Lyons, J.I., 1988. Volcanogenic iron oxide deposits, Cerro deMercado and vicinity, Durango,Mexico. Econ. Geol. 83, 1886–1906.

Ma, F., Jiang, S.Y., Jiang, Y.H., Wang, R.C., Ling, H.F., Ni, P., 2006. Pb isotope research ofporphyrite Fe deposits in the Ningwu Area. Acta Geol. Sin. 80, 279–286 (in Chinesewith English abstract).

Mao, J.W., Wang, Y.T., Lehmann, B., Yu, J.J., Du, A.D., Mei, Y.X., Li, Y.F., Zang, W.S., Stein, H.J.,Zhou, T.F., 2006. Molybdenite Re–Os and albite 40Ar/39Ar dating of Cu–Au–Mo andmagnetite porphyry systems in the Yangtze River valley andmetallogenic implications.Ore Geol. Rev. 29, 307–324.

Markl, G., 2001. A new type of silicate liquid immiscibility in peralkaline nepheline syenites(lujavrites), the Ilímaussaq complex, South Greenland. Contrib. Mineral. Petrol. 141,458–472.

Martinsson, O., 1997. Tectonic setting andmetallogeny of the Kiruna greenstones. Ph.D.thesis, Luleå University of Technology, Luleå, Sweden.

Martinsson, O., Weihed, P., 1999. Metallogeny of juvenile palaeoproterozoic volcanicarcs and greenstone belts in rifted Achaean crust in the northern part of Sweden.In: Stanley, C.J., et al. (Ed.), Fennoscandian Shield: Mineral Deposits-Processes toProcessing, vol. 2, pp. 1329–1332.

Naslund, H.R., 1983. The effects of oxygen fugacity on liquid immiscibility in iron-bearingsilicate melts. Am. J. Sci. 283, 1034–1059.

Naslund, H.R., Aguirre, R., Dobbs, F.M., Henriquez, F.J., Nyström, J.O., 2000. The origin,emplacement, and eruption of ore magmas. IX Congreso Geologico Chileno, vol.2. Sociedad geológica de Chile, pp. 135–139.

Naslund,H.R.,Henriquez, F., Nyström, J.O., Vivallo,W., Dobbs, F.M., 2002.Magmatic iron oresand associated mineralization: examples from the Chilean high Andes and coastal cor-dillera. In: Porter, T.M. (Ed.), Hydrothermal iron oxide copper-gold and related deposits—a global perspective, vol. 2. PGC Publishing, Adelaide, pp. 207–226.

Ningwu Research Group, 1978. Ningwu Porphyry Iron Ores. Geological PublishingHouse, Beijing. (in Chinese).

Nyström, J.O., Henriquez, F., 1994. Magmatic features of iron ores of the Kiruna type inChile and Sweden: ore textures andmagnetite geochemistry. Econ. Geol. 89, 820–839.

Pan, Y.M., Dong, P., 1999. The Lower Changjiang (Yangzi/Yangtze River) metallogenicbelt, east central China: intrusion- and wall rock-hosted Cu–Fe–Au, Mo, Zn, Pb,Ag deposits. Ore Geol. Rev. 15, 177–242.

Parák, T., 1975. Kiruna iron ores are not “intrusive-magmatic ores of the Kiruna type”.Econ. Geol. 70, 1242–1258.

Park Jr., C.F., 1961. A magnetite “flow” in northern Chile. Econ. Geol. 56, 431–436.Philpotts, A.R., 1967. Origin of certain iron titanium oxide and apatite rocks. Econ. Geol.

62, 303–316.Ren, Q.J., 1991. Studies of Ti–Fe oxides in igneous rocks and related deposits. Science

Press, Beijing. (in Chinese).Ripa,M., 1988. Geochemistry of wall-rock alteration and ofmixed volcanic exhalative facies

at the Proterozoic Stollberg iron–lead–zinc–manganese–silver deposit, Bergslagen,Sweden. Geol. Mijnbouw 67, 443–457.

Roedder, E., 1979. Silicate liquid immiscibility inmagmas. In: Yoder,H.S.J. (Ed.), The Evolutionof the Igneous Rocks, 50th Anniversary Perspectives. Princeton University Press,pp. 15–57.

Roedder, E., 1984. Fluid inclusions. Reviews in Mineralogy, 12. Mineralogical Society ofAmerica, Washington, D.C.

Ryerson, F.J., Hess, P.C., 1980. The role of P2O5 in silicate melts. Geochim. Cosmochim.Acta 44, 611–624.

Sillitoe, R.H., Burrows, D.R., 2002. New field evidence bearing on the origin of the ElLaco magnetite deposit, Northern Chile. Econ. Geol. 97, 1101–1109.

Song, X.X., Chen, Y.C., Sheng, J.F., Ai, Y.D., 1981. On iron deposits formed fromvolcanogenic–hypabyssal ore magma. Acta Geol. Sin. 1, 41–55 (in Chinese with English abstract).

Sørensen, H., Bailey, J.C., Kogarko, L.N., Rose-Hansen, J., Møller, S.K., 2003. Spheroidalstructures in arfvedsonite lujavrite, Ilímaussaq alkaline complex, South Greenland—an example of macro-scale liquid immiscibility. Lithos 70, 1–20.

Streck, M.J., 2008. Mineral textures and zoning as evidence for open system processes.Rev. Mineral. Geochem. 69, 595–622.

Suk, N.I., 1998. Distribution of ore elements between immiscible liquids in silicate–phosphatesystems (experimental investigation). Acta Univ. Carol. Geol. 42, 138–140.

Tang, Y.C.,Wu, C.Y., Chu, G.Z., Xing, F.M.,Wang, Y.M., Cao, F.Y., Chang, Y.F., 1998. Geology ofCopper–Gold Polymetallic Deposits in the Along-Changjiang Area of Anhui Province.Geological Publishing House, Beijing. (in Chinese with English abstract).

Tepley, F.J., Davidson, J., Tilling, R.I., Arth, J.G., 2000. Magma mixing, recharge and eruptivehistories recorded in plagioclase phenocrysts fromEl ChichónVolcano,Mexico. J. Petrol.41, 1397–1411.

Thompson, A.B., Aerts, M., Hack, A.C., 2007. Liquid immiscibility in silicate melts and relatedsystems. Rev. Mineral. Geochem. 65, 99–127.

Tsuchiyama, A., Takahashi, E., 1983. Melting kinetics of a plagioclase feldspar. Contrib.Mineral. Petrol. 84, 345–354.

Turner, M.B., Cronin, S.J., Smith, I.E., Stewart, R.B., Neall, V.E., 2008. J. Volcanol.Geotherm. Res. 177, 1063–1076.

Veksler, I.V., 2009. Extreme iron enrichment and liquid immiscibility in mafic intrusions:experimental evidence revisited. Lithos 111, 72–82.

Veksler, I.V., Dorfman, A.M., Danyushevsky, L.V., Jakobsen, J.K., Dingwell, D.B., 2006. Im-miscible silicate liquid partition coefficients: implications for crystal-melt elementpartitioning and basalt petrogenesis. Contrib. Mineral. Petrol. 152, 685–702.

Visser, W., Koster van Groos, A.F., 1979a. Effect of pressure on liquid immiscibility inthe system K2O–FeO–Al2O3–SiO2–P2O5. Am. J. Sci. 279, 1160–1175.

Visser,W., Koster van Groos, A.F., 1979b. Effects of P2O5 and TiO2 on liquid–liquid equilibriain the system K2O–FeO–Al2O3–SiO2. Am. J. Sci. 279, 970–988.

Visser, W., Koster van Groos, A.F., 1979c. Phase relations in the system K2O–FeO–Al2O3–SiO2 at 1 atmosphere with special emphasis on low temperature liquid im-miscibility. Am. J. Sci. 279, 70–91.

Webster, J.D., Tappen, C.M., Mandeville, C.W., 2009. Partitioning behavior of chlorineand fluorine in the system apatite–melt–fluid. II: Felsic silicate systems at 200MPa. Geochim. Cosmochim. Acta 73, 559–581.

Xing, F.M., 1998. Deep tectonic information of magmatic rocks along the Yangtze Riverin Anhui. Reg. Geol. China 17, 195–200 (in Chinese with English abstract).

Xu, Z.G., 1990. Mesozoic volcanism and volcanogenic iron-ore deposits in easternChina. Special Paper, 237. Geological Society of America, pp. 32–34.

Xu, X., Xing, F.M., 1994.Whole-rock andmineral Rb–Sr isochron ages of the three gabbrosin Nanjing-Wuhu area, China. Sci. Geol. Sin. 29 (3), 309–312 (in Chinese with Englishabstract).

Yan, J., Liu, H.Q., Song, C.Z., Xu, X.S., An, Y.J., Dai, L.Q., 2009. Zircon U–Pb geochronologyof the volcanic rocks from Fanchang–Ningwu volcanic basins in the Lower Yangtzeregion and its geological implications. Chinese Science Bulletin 54, 1716–1724.

Yu, J.J., Mao, J.W., 2002. Rare elements in apatite from porphyry iron deposits ofNingwu area. Mineral Deposit 21, 65–73 (in Chinese with English abstract).

Yu, J.J., Mao, J.W., Zhang, C.Q., 2008. The possible contribution of a mantle-derived fluid tothe Ningwu porphyry iron deposits—evidence from carbon and strontium isotopes ofapatites. Progress in Natural Science 18, 167–172.

Yu, J.J., Chen, Y.C., Mao, J.W., Pirajno, F., Duan, C., this issue. Review of geology, alterationand origin of iron oxide-apatite deposits in the Cretaceous Ningwu basin, Lower

346 T. Hou et al. / Ore Geology Reviews 43 (2011) 333–346

YangtzeRiverValley, easternChina: implications for ore genesis andgeodynamic setting,Ore Geology Reviews doi:10.1016/j.oregeorev.2011.07.008.

Zhai, Y.S., Yao, S.Z., Lin, X.D., Zhou, X.R., Wan, T.F., Jin, F.Q., Zhou, Z.G., 1992. MetallogenicRegularities of Iron–Copper (Gold) Deposits in Middle–Lower Yangtze River Valley.Geological Publishing House, Beijing. (in Chinese).

Zhai, Y.S., Xiong, Y.L., Yao, S.Z., Lin, X.D., 1996. Metallogeny of copper and iron depositsin the Eastern Yangtse Craton, east-central China. Ore Geol. Rev. 11, 229–248.

Zhang, Q., Jian, P., Liu, D.Y.,Wang, Y.L., Qian, Q., Xue, H.M., 2003. SHRIMPdating of volcanicrocks from Ningwu area and geological implication. Science in China (Series D) 33,830–837 (in Chinese with English abstract).

Zhang, Z.C., Mao, J.W., Saunders, A.D., Ai, Y., Li, Y., Zhao, L., 2009. Petrogenetic modeling ofthree mafic–ultramafic layered intrusions in the Emeishan large igneous province,SW China, based on isotopic and bulk chemical constraints. Lithos 113, 369–392.

Zhao, Y.C., 1993a. Geological characteristics and using prospects of associated apatites oreswithin the porphyritic iron deposit in Nangjing-Wuhu area. Geology and Prospecting29 (12), 6–12 (in Chinese with English abstract).

Zhao, Y.X., 1993b. Mechanism of Formation of the Contact Iron Deposits along the Middle–Lower Reaches of the Yangtze River. China University of Geosciences Press,Wuhan. (inChinese with English abstract).

Zhou,M.F., Robinson, P.T., Lesher, C.M., Keays, R.R., Zhang, C.J.,Malpas, J., 2005. Geochemistry,petrogenesis and metallogenesis of the Panzhihua gabbroic layered intrusion and asso-ciated Fe–Ti–V oxide deposits, Sichuan Province, SW China. J. Petrol. 46, 2253–2280.