Lithos 156–159 (2013) 218–229

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The Beiminghe skarn iron deposit, eastern China: Geochronology, isotopegeochemistry and implications for the destruction of the North China Craton

Jun-Feng Shen a,⁎, M. Santosh a,b, Sheng-Rong Li a, Hua-Feng Zhang a, Na Yin a, Guo-Cheng Dong a,Yan-Juan Wang a, Guang-Gang Ma a, Hong-Jun Yu a

a School of Earth Science and Resources, China University of Geosciences, Beijing 100083, Chinab Division of Interdisciplinary Science, Faculty of Science, Kochi University, Kochi 780-8520, Japan

⁎ Corresponding author. Tel.: +86 10 8232 1732.E-mail address: shenjf@cugb.edu.cn (J.-F. Shen).

0024-4937/$ – see front matter © 2012 Elsevier B.V. Allhttp://dx.doi.org/10.1016/j.lithos.2012.11.003

a b s t r a c t

a r t i c l e i n f o

Article history:Received 29 May 2012Accepted 6 November 2012Available online 13 November 2012

Keywords:Skarn iron depositGeochronologyIsotope geochemistrySouth Taihang MountainNorth China Craton

The Beiminghe (BMH) iron ore in the southern part of the Taihang Mountain (TM), Hebei province, is one of thelargest skarn iron deposits in China. Here we report phlogopite 40Ar–39Ar and zircon U–Pb age data, as well assulfur, lead, and He–Ar isotope geochemistry of pyrite from the ores and skarnitized rocks in the deposit in anattempt to constrain the timing and mechanism of formation of the mineralization. The phlogopite 40Ar–39Arand LA-ICP-MS zircon U–Pb data show markedly consistent ages constraining the timing of ore formation as136–137 Ma. The presence of several inherited zircons with late Archean or Paleoproterozoic ages indicatesthe participation of the basement rocks during the ore-forming process. The δ34S values of pyrite from theores range from 12.2 to 16.5‰, with 206Pb/204Pb=17.84–18.79, 207Pb/204Pb=15.46–15.62, and 208Pb/204Pb=37.93–39.75, suggesting that continental crust is the major contributor. This is further confirmed by the He–Arisotope data (3He/4He=0.0648–0.1886 Ra, mean 0.1237Ra; 40Ar/36Ar=311.7–22909.4; and 40Ar⁎/4He=0.036–0.421). The Mesozoic magmatism and metallogeny in the BMH correlate well with the peak event oflithospheric thinning and destruction of the North China Craton during this process, the early Precambrianlower crustal rocks in the region were re-melted through underplating of mantle magmas, leading to the forma-tion of the Beiminghe monzodioritic pluton. Minor mantle input occurred during the evolution of themonzodiorite magma, which scavenged the ore-forming materials from the lower crust. Interaction of themagmas and fluids with the surrounding rocks resulted in the formation of the Beiminghe skarn iron deposits.The magmatism and metallogeny in the Taihang Mountain are signatures of the extensive craton destructionand lithospheric thinning in the eastern part of the North China Craton during Mesozoic, probably associatedwith Pacific slab subduction.

© 2012 Elsevier B.V. All rights reserved.

1. Introduction

The Beiminghe (BMH) iron deposit is located at the southern part ofthe Taihang Mountain (TM), ca. 10 km from northwest of the Wu'anCity in the Hebei province of China. Nearly one hundred skarn-typeiron deposits termed as “mineralizing districts” with comparable fea-tures have been recognized, which are described in Chinese literatureas the Handan–Xingtai or Han-Xing skarn iron deposits (Feng, 1998;Zheng et al., 2007a, 2007b, 2007c). The Han-Xing skarn iron depositshave been widely accepted as products of contact metamorphismbetween the Mesozoic pluton and the surrounding Ordovician carbon-ate sedimentary strata (Feng, 1998; Niu et al., 1995; Zhang et al.,2009; Zheng et al., 2007a, 2007b, 2007c). Although some previous stud-ies (Chen et al., 2005a; Zheng et al., 2007a) have attempted to proposethe ore-forming model, no detailed information is available on thetiming and mechanism of ore formation in this area.

rights reserved.

In this study, we present new geochronological data and mineralisotopic compositions of pyrites from ores and country rocks from theBMH iron deposit. Based on the results, we evaluate the age of oreformation, the source of ore-forming materials, and the geodynamicsetting of ore genesis.

2. Geological setting

2.1. Regional geology

The major lithological units of the TM are composed of early Pre-cambrian metamorphic rocks, Phanerozoic sedimentary units andMesozoic magmatic intrusions (Li et al., 2012). The basement mainlyconsists of TTG (tonalite–trondhjemite–granodiorite) gneisses and am-phibolites, the protoliths of which formed during Meso-Neoarchean,andwere latermetamorphosed during late Paleoproterozoic, associatedwith the final cratonization of the NCC (Geng et al., 2012; Li et al., 2012;Liu et al., 2011; Zhai and Santosh, 2011; Zhao et al., 2000). The skarniron deposits of MBH are located within the eastern periphery of the

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Central Orogenic Belt, also termed as the Trans-North China Orogen(TNCO) adjacent to the western margin of the Eastern Block of theNCC (Fig. 1A,B, Xu et al., 2009b, 2009c; Zhai and Santosh, 2011; Zhaoet al., 2001; Zhu et al., 2011). Regionally, the sedimentary rocks in thisarea are dominated by Cambrian–Ordovician carbonates in thewesternpart, and Carboniferous–Permian clastic rocks in the eastern part(Fig. 2).

The Mesozoic magmatic rocks intruding the Precambrian basementas well as the younger sedimentary cover belong to a major NNE-trending magmatic belt (Li et al., 2012). Several batholiths, stocks andsills mainly composed of intermediate calc-alkaline rocks have beenidentified (Chen et al., 2006; Li et al., 2012; Liu et al., 2009; Peng et al.,2004; Zheng et al., 2007a, 2007b; Zhou and Chen, 2005).

2.2. Ore geology

The BMH skarn iron deposit belongs to the Wu'an iron clusterregion, composed of several iron deposits to the west of the Wu'ancity, surrounding the major pluton (Fig. 2). The dominant basementrocks in this region are the Archean Zanhuang Complex, intruded byMesozoic intermediate plutons comprising mostly of monzonite,monzodiorite, diorite, and quartz–diorite (Chen et al., 2009). The gene-sis of these intrusions has been linked to the subduction of the Paleo-Pacific plate (Zheng et al., 2007c). The major sedimentary unit in themine is the Majiagou limestone, deposited during Middle Ordovician.

The main ore-controlling structure in the mine is an anticlinetrending NW. The ore bodies typically occur at the contact zones ofthe dioritic intrusion and the Majiagou limestone. From the contactzone profile, the shape of the ore bodies is reconstructed as complexlenses with end-to-end discontinuity. The serrated and interspersed

Fig. 1. The location of Beiminghe iron skarn deposit at the eastern periphery of the Trans-NCraton. WB — Western Block. EB — Eastern lock. TNCO — Trans-North China Orogen.Modified after Xu et al. (2009a), Zhai and Santosh (2011) and Zhao et al. (2001).

features of the ore bodies are visible near their edges, with both inwardand outward contact zones. Field studies reveal a close relationship be-tween the diorite, which at some places grades to diorite porphyrite ormonzodiorite, and the formation of the skarn and ores. We show inFig. 3 a horizontal level cross section at 110 m below the surface.

Alteration zones are clearly developed at the contact of the intrusivewith the surrounding limestone, and these zones can be divided into al-tered diorite zone, the endoskarn zone, the sulfide–magnetite zone, theexoskarn zone and the marble zone. In general, the endoskarn zoneshows a wider distribution as compared to the exoskarn zone whichis narrow. The diopside-type skarn dominates the BMH iron deposit,with occasional garnet-type. The skarn mineral compositions basicallyinclude diopside, tremolite, actinolite, phlogopite, humite, serpentine,garnet, and epidote. The marble at the external contact zone often hasbeen brecciated or altered by chlorite. The main metallic minerals aremagnetite, pyrite, and a small amount of hematite, chalcopyrite,pyrrhotine, and nickeliferous magnetite. The content of magnetite inthe ores varies from 50 to 80%. The subhedral to anhedral coarse mag-netite grains range from 0.15 to 0.4 mm in size. Under the influence oflate hydrothermal alteration, the magnetite grains are replaced by py-rite, calcite, and hematite, with the relict texture preserved in manycases. The pyrite is subhedral to idiomorphic and generally ranges incontent from 1 to 10%, and sometimes up to 25–50% in some domains.

3. Sampling and analytical methods

3.1. Sampling

The phlogopite-bearing sample for 40Ar–39Ar dating was collectedfrom the exoskarn zone at a depth of −230 m in the mine. Here,

orth China Orogen. The inset figure shows the major tectonic units of the North China

Fig. 2. Geological map and distribution of iron deposits in Handan district, Hebei province.After Zheng et al. (2007a).

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grey-green phlogopite-rich vein occurs in the skarn (Fig. 4a), and themineral shows intergrowth with magnetite (Fig. 4b).

Two representative samples of the magmatic intrusive were select-ed for zircon separation and U–Pb dating: one from 110 m depth (sam-ple B110-9-2), and the other from 245 m depth (sample B245-1).

Sample B110-9-2 was collected from weakly altered dioriteporphyrite, away from the ore body in the main tunnel. The sampleexhibits porphyritic texture with phenocrysts of euhedral plagioclase(Fig. 5a) and minor hornblende. Taxitic and striped texture are alsovisible at places. The plagioclase shows marginal albitization andsericitic alteration. Albite also occurs as coarse tabular and columnarcrystals. The major mineral assemblage in this sample is plagioclase(ca. 90%) and hornblende (ca. 10%), with occasional quartz and calciteveins (Fig. 5b).

Sample B245-1 was taken from the skarnitized diorite close to theore body (Fig. 5c). The sample is intensely altered, with hydrothermalcalcite veins visible in the hand specimens, and some of the domainsare cracked and brecciated. The breccias are grey to black with grains6–8 mm across set within ferruginous cement. Under the microscope,the intense alteration zone shows abundant diopside. The alternation

Fig. 3. The horizontal level cross section g

assemblage can be divided into two stages: stage 1 where diopsideoccurs as small grains and stage 2 where it occurs as euhedral coarsegrains (Fig. 5d). Calcite veins are partly developed, occasionally inassociation with disseminated pyrite and chalcopyrite.

3.2. Analytical methods

3.2.1. 40Ar/39Ar datingThe skarn containing phlogopite grains were crushed and sieved,

and purified phlogopite sample was obtained using the conventionaltechniques of magnetic separators and heavy liquids. The phlogopitefor 40Ar/39Ar analysis was purified to 99%, and then ultrasonicallycleaned in distilled water and ethanol.

The 40Ar/39Ar dating of phlogopite was carried out using stepwiseincremental heatingmethod at the Institute of Geology, Chinese Acade-my of Geological Sciences (CAGS). Instrumental conditions and analyti-cal details are as described by Chen et al. (2002). The purified sampleswere wrapped in aluminum foil and loaded into a tube of aluminumfoil. Each tube contains 2–3 monitors (an internal standard: Fangshanbiotite, 132.7±1.2 Ma, Chen et al., 2002) in between the minerals. A

eological map at height of −110 m.

Fig. 4. Field and thin section photographs of phlogopite and associated rocks and minerals from the Beiminghe ore deposit. a. Phlogopite in the exoskarn zone, from the mine tunnel.b. Phlogopite associated with diopside and magnetite. The phlogopites were collected for 40Ar–39Ar dating.

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number of such tubes were sealed into quartz vial and irradiated for2941 min in the nuclear reactor at the Chinese Academy of Atomic En-ergy. The reactor delivers a neutron flux of about 6.0×1012n cm−2 s−1;the integrated neutron flux is about 1.15×1018n cm−2. The irradiatedsamples and monitors were loaded into the vacuum extraction systemand baked for 48 h at 120–150 °C. The Ar extraction system comprisesan electron bombardment heated furnace in which the samples areheated under vacuum. The duration is 30 min for heating–extractionat each temperature increment, and 30 min for purification. The puri-fied Ar was trapped in activated charcoal finger at liquid–nitrogen tem-perature, and then released into the MM-1200B Mass Spectrometer to

Fig. 5. Field photograph and photomicrographs of the samples collected for zircon U–Pb dating.(crossed nicols) of the sample. c. Location of skarnitized diorite sample B245-1 inside mine tunPl— plagioclase; Di— diopside; Cal — calcite.

analyze Ar isotopic ratios. Measured isotopic ratios were corrected formass discrimination, atmospheric Ar component, procedural blanksand mass interference induced by irradiation. The blanks of m/e of40Ar, 39Ar, 37Ar and 36Ar are less than 6×10−15, 4×10−16, 8×10−17

and 2×10−17 mol, respectively. The correction factors of interferingisotopes produced during irradiation were determined by the analysisof irradiated K2SO4 and CaF4 pure salts. The values are: (40Ar/39Ar)k=0.004782; (36Ar/37Ar)Ca=0.000240; and (39Ar/37Ar)Ca=0.000806.Ages were calculated using the ISOPLOT program (version 2.49,Ludwig, 2001). The K decay constant used was 0.5543 Ga−1. All 37Arwere corrected for radiogenic decay (half-life 35.1 days). Uncertainties

a. Location of diorite porphyrite sample B 110-9-2 insidemine tunnel. b. Photomicrographnel. d. Photomicrograph (parallel nicols) of the sample. Abbreviations. Hbl— hornblende;

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on the apparent ages on each step are quoted at the 1σ level, butweightedmean plateau ages and isochron ages are given at the 2σ level.

3.2.2. Zircon U–Pb datingSeparation of zircon grains was performed using the conventional

techniques including heavy liquids, magnetic separator and hand-picking under a binocular microscope fitted with a UV light. Zircongrains were mounted in epoxy disks, polished to expose thehalf-sections of grains, and then coated with gold.

Prior to the U–Th–Pb dating, the internal textures of the zirconswere studied by transmitted and reflected lights microscope, andcathodoluminescence (CL) technique using a CAMECA SX51 at theKey Laboratory of Metallogeny and Mineral Assessment, Institute ofMineral Resources, CAGS. The procedures are the same as those de-scribed by Hou et al. (2009).

LA-ICP-MS isotope U–Pb dating was performed by the conventionalisotope dilution technique of single zircons after air abrasion. Fractionsize varied somewhat depending on U content and grain size of zircon.Zircon digestion, separation of U and Pb, and isotopedilutionmass spec-trometry using a 205Pb–235U enriched tracer solution were madefollowing the procedure described by Krogh (1982). The total proce-dure blanks for Pb and U are less than 0.05 and 0.002 ng, respectively.Isotope ratios are measured by single collector peak jumping VG-354mass spectrometer equipped with a Daly-type detector operating inion-counting mode at Tianjin Institute of Geology and MineralResources. Errors to the atomic and isotopic ratios are quoted at the2σ confidence level. Common Pb corrections are made by using themodel Pb composition of Stacey and Kramers (1975). The age calcula-tions were performed by means of the ISOPLOT program of Ludwig(1994).

Table 1Ar isotopic data for phlogopite in BMH iron deposit.

T (°C) (40Ar/39Ar)m (36Ar/39Ar)m (37Ar/39Ar)m (38Ar/39Ar)m 40Ar(%)

700 57.0512 0.1848 0.459 0.1766 4.35800 117.0246 0.3551 0.0957 0.1247 10.33860 53.1192 0.1331 0 0.0504 25.97940 46.1235 0.1081 0.0501 0.0451 30.77980 16.4144 0.0123 0 0.0023 77.831020 16.1043 0.0121 0.0156 0.023 77.761060 16.0667 0.0119 0 0.0232 78.081100 15.4623 0.0112 0 0.0222 78.581160 16.3979 0.0147 0.0082 0.0223 73.481220 18.4316 0.0225 0 0.0242 63.841400 49.714 0.0853 0 0.0554 49.26

F 39Ar (×10−14 mol) 39Ar (Cum.) (%) Age (Ma) ±1 Ma

2.4817 0.42 2.52 27.8 212.0907 0.32 4.43 131.7 513.794 0.43 6.99 149.5 2.414.1919 2.65 22.87 153.6 1.512.7754 2.69 39.05 138.9 1.412.5229 3.91 62.51 136.2 1.312.5442 2.35 76.61 136.4 1.312.1501 1.85 87.75 132.3 1.312.0489 1.68 97.83 131.2 1.311.7674 0.3 99.62 128.3 2.624.4915 0.06 100 257.4 7.7

Note: W=30.56 mg, J=0.006263, F=40Ar⁎/39Ar, is the ratio of radiogenic Argon40and Argon39.

3.2.3. S/Pb/He–Ar isotopic compositionThe sample preparation for sulfur isotope analyses followed the pro-

cedures outlined by Glesemann et al. (1994) and themeasurementwasperformed on a MAT-251EM mass spectrometer at the Stable IsotopeLaboratory of CAGS. Data are reportedwith an accuracy of ±0.2‰ (2σ).

The Pb isotope analysis was performed at the CAGS, and the proce-dure adopted is as follows. The whole-rock Pb was separated by anionexchange onHCl–Br columns.Within the analytical period, 30measure-ments of NBS981 gave average values of 206Pb/204Pb=16.937±1 (1σ),207Pb/204Pb=15.491±1, and 208Pb/204Pb=36.696±1. The BCR-2standard gave 206Pb/204Pb=18.742±1 (1σ), 207Pb/204Pb=15.620±1, and 208Pb/204Pb=38.705±1. Total procedural Pb blanks were inthe range of 0.1–0.3 ng.

He–Ar gas isotope analyses were performed with a MM5400 gasesmass spectrometer (Micromass, GB) at Lanzhou Center for Oil and GasResources, Institute of Geology and Geophysics, China Academy ofSciences. Experiment was done at electric current It4=800 μA, It40=200 μA, and high voltage 9.000 kV. All weighed samples of pyrite foranalysis were packed into aluminum foil and shifted to the cruciblefor gas extraction under high vacuum conditions. When a pressurelower than 1×10−5 Pa was attained, the samples were heated at130 °C for at least 10 h to eliminate secondary fluid inclusions andtrace gases occurring in cleavages or fractures in the crusts. The sampleswere then fused at high temperatures of up to 1600 °C, and the releasedgases were purified through activated charcoal traps at the liquid nitro-gen temperature to separate He and Ar fromNe+Kr+Xe for He and Aranalyses on the mass spectrometer, respectively. The minimum heatblanks for the MM5400 mass spectrometer at 1600 °C are: 4He=1.10×10−14 mol; 20Ne=1.82×10−14 mol; 40Ar=6.21×10−13 mol;84Kr=1.37×10−16 mol; and 132Xe=5.65×10−18 mol. The standardfor normalizing the analytical results is air in Lanzhou (AIRLZ2003). De-tailed sample preparation andmeasurement procedures followed thosein He et al. (2011) and Ye et al. (2001, 2007).

4. Results

4.1. 39Ar–40Ar age

The stepwise incremental heatingmethod depends on the apparentage to derive credible data (Chen et al., 2005b). If the 39Ar release is over50% during more than three heating steps, and the plateau age of eachstep is less than 2σ, the corresponding plateau age is considered asauthentic (Dalrymple and Lanphere, 1974). Due to argon loss inducedby mineral margin or lattice defects, abnormal values could result ineither low or high temperature step. Therefore, the medium–high tem-perature step most likely represents the real age (Chen et al., 2005b).

As shown in Table 1 and Fig. 6, the sample heating procedure isdivided into 10 steps (700–1400 °C). An unstable plateau appears atthe initial low temperature (700–940 °C), with apparent age rangingfrom 27.8 to 153.6 Ma. However, the deviation of plateau data calculat-ed from 700 °C to 860 °C is over 2σ, with low 39Ar 22.86%. These couldbe initiated by little argon loss of mineral margin (Chen et al., 2005b),and may have no relation with the actual ore-forming age.

Over 1100 °C, the 39Ar release clearly decreases. Although there is afluctuation from 1220 to 1400 °C, considering the deviation of plateaudata to be over 2σ, the data might not represent the true crystallizationage of phlogopite. Only plateau data calculated from the 980 to 1060 °Cstep tends to be continuous and steady, and represents 53% of the 39Arreleased. The plateau age data calculated from this thermal step shows137±2 Ma, which we consider to represent the crystallization age ofphlogopite in this ore deposit.

Furthermore, ignoring the two abnormal values (27.8 Ma at 700 °Cand 257.4 Ma at 1400 °C, respectively), the correlation of other 9 datais R2=0.997,with amean age of 137.5 Ma,which is close to the plateauage 137±2 Ma obtained from the 980 to 1060 °C step. From the corre-lation of 40Ar/36Ar and 39Ar/36Ar, we obtain the initial value of 40Ar/36Arfor sample at 298.6, which is similar to the standard value of atmo-spheric argon 295.5. This further confirms that the age of 137 Ma reli-ably represents the crystallization age of phlogopite.

Since the field studies and petrography clearly indicate a cogeneticnature of the phlogopite and magnetite, the age obtained from phlogo-pite can be taken as the timing of mineralization in the BMH irondeposit.

Fig. 6. Plateau age plot of phlogopite in BMH iron deposit.

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4.2. Zircon U–Pb age

4.2.1. Zircon CL imageMost zircon grains from the diorite porphyrite are short columnar or

anhedral ellipsoidal grains. Some grains show anhedral morphology,whereas few others are well-formed idiomorphic crystals. As seenfrom Fig. 7 many of the zircons display clear core–rim structure. Someof the grains are cracked and contain fractures.

Zircon grains in our sample can be divided into two types in terms ofmorphology and internal structure. The first type is inherited zircons,mostly forming the core domains and is white or grayish under theCL. Some of these grains show clear zoning. Based on the presence orabsence of zoning, the inherited zircons can be further divided intomagmatic zircons with bands and structure metamorphic zircons.Among these, some grains show intense etching and inhomogeneousoptical characteristics.

The second type is neo-formed zircons, mostly developed in the rimof grains, with various core–rim structures, indicating crystallizationfrom magma.

4.2.2. Zircon U–Pb ageThe zircon U–Pb analytic data and computed ages are reported in

Tables 2 and 3.The data for sample B110-9-2 (Table 2) show considerable spread

and the 17 analyses can be divided into 4 groups as described below.The first group includes 4 analyses with ages ranging from 134 to

137 Ma, and defines a concordant age of 136±2 Ma (Fig. 8a). Amongthese, points 14 and 19 are from domains showing magmatic texture.In particular, point 19 represents newly formed magmatic zircon.

Fig. 7. CL images of zircons from the diorite porphyrite. The analytical spots and agedata are also shown.

The second group consists of 5 data with ages ranging from 2500 to2786 Ma. These belong to the inherited zircons with clear dissolutionfeatures and blurred growth edges.We consider these ages to representthe inherited zircons, derived from a Neoarcheanmagmatic protolith inthe basement.

The third group is defined by 5 data, with ages ranging from 2216 to2408 Ma. This group is identical to the previous group in morphologyand internal structure, and represents inherited zircons from aPaleoproterozoic magmatic source.

The fourth group, represented by only 2 results, defines ages of 1553and 1898 Ma. These spots are from structureless inherited domains andtherefore represent Paleo- to Mesoproterozoic metamorphic zircons.

The analytical data from sample B245-1 shown in Table 3 spreadalong the Concordia,with only fewdeviations. The data define a concor-dant age at 2482±20 Ma (Fig. 8b). The majority of data defineNeoarchean and Paleoproterozoic ages, clearly indicating closely com-parable with those from sample B110-9-2.

4.3. Mineral isotopic compositions

4.3.1. Sulfur isotopePrevious studies (Liu and Shi, 1998) have reported δ34S from the

skarn iron ore in Han-Xing which shows a range of 7.27 to 16.36‰,values that are higher than the range of meteorite S. In the presentstudy, the δ34S values of pyrite from the BMH iron deposit show arange of 12.2 to 16.5‰ (Table 4). Although the range is relativelynarrow, the values do not correspond with that of meteorite S isotope(Fig. 9, where the data from Beiminghe iron are plotted together withthose reported by Zhang et al., 2009 and Zheng, 2007d). Thus, theδ34S values of the BMH iron deposit do not provide any clear mantle Ssignature, and instead indicate the involvement of crustal materials inthe metallogenic process.

4.3.2. Lead isotopic compositionFive pyrite samples analyzed in this study (Fig. 10), show 206Pb/

204Pb values of 17.84–18.79 (average 18.42); 207Pb/204Pb values of15.46–15.62 (average 15.56) and 208Pb/204Pb values of 37.93–39.75(average 38.73). In the evolution models of 207Pb/204Pb with 206Pb/204Pb (Fig. 10a), five analyses fall between the orogenic belt and themantle evolution lines, mainly clustering near the orogenic belt line.207Pb/204Pb and 206Pb/204Pb values of intrusive rocks in southernTM plot near the lower crust and mantle evolution lines (coloredregion in Fig. 10b; additional data from Cai et al., 2004a and Zhanget al., 2009).

In the evolutionmodel of 208Pb/204Pb and 206Pb/204Pb, the analyticaldata also plot near the orogenic belt line trending towards the field oflower crust; this is further supported by similar values from the intru-sive rocks in southern TM (colored region Fig. 10b; additional datafrom Zhang et al., 2009).

4.3.3. He–Ar isotopic compositionsNumerous studies reveal that the He and Ar gases trace the charac-

teristics of differentiation alongwith the degassing in the Earth's evolu-tion (Baptiste and Fougute, 1996; Barnard et al., 1994a, 1994b; Dunaiand Touret, 1995; Hu, 1997; Hu et al., 1999; Li et al., 2004; Marty etal., 1989; Matthews et al., 1987; Simmons et al., 1987; Stuart et al.,1995; Wu et al., 2003). Thus, the He and Ar isotopes are distinct forthe mantle, crust and surface atmosphere. Among these, the mantleretains higher original He and Ar (such as 3He and 36Ar), but lowerradiogenic 4He and 40Ar. Therefore, it is generally considered that thevalue of 3He/4He in the continental mantle is 6–9 Ra (Ra=3He/4He=1.400×10−6), the value of 40Ar/36Ar is greater than 20,000, and thevalue of 40Ar/4He is about 0.33–0.56 (Cai et al., 2004b; Hu et al., 1999;Simmons et al., 1987). However, the value of 3He/4He in the typicalcrust is usually less than 0.1 Ra, in most cases even only 0.01–0.05 Ra.As there is abundant radiogenic 40Ar in the crust, the concentration of

Table 2Results of zircon LA-ICP-MS dating on sample B110-9-2.

Spot U(ppm)

Th(ppm)

232Th/238U

Error 206Pb/238U

Error 207Pb/235U

Error 207Pb/206Pb

Error 206Pb/238U(age)

Error 207Pb/235U(age)

Error 207Pb/206Pb(age)

Error

1 73 14 1.4384 ±0.0152 0.5229 ±0.0043 13.3658 ±1.0474 0.1854 ±0.0126 2711 ±22 2706 ±212 2702 ±1832 47 8 0.5441 ±0.0052 0.5096 ±0.0102 11.3609 ±2.0362 0.1617 ±0.0302 2655 ±53 2553 ±458 2474 ±4635 234 36 0.547 ±0.0034 0.4437 ±0.0030 9.5144 ±0.4078 0.1555 ±0.0066 2367 ±16 2389 ±102 2407 ±1036 206 46 0.2144 ±0.0019 0.0274 ±0.0004 0.8476 ±0.1354 0.2245 ±0.0471 174 ±2 623 ±100 3013 ±6327 83 14 0.6095 ±0.0039 0.4777 ±0.0040 10.7462 ±0.7235 0.1631 ±0.0107 2517 ±21 2501 ±168 2489 ±1648 480 23 0.4535 ±0.0017 0.0214 ±0.0002 0.1436 ±0.0091 0.0486 ±0.0031 137 ±1 136 ±9 128 ±810 75 12 0.2739 ±0.0009 0.4497 ±0.0035 9.7175 ±0.6120 0.1567 ±0.0098 2394 ±19 2408 ±152 2421 ±15111 126 19 0.4744 ±0.0053 0.2685 ±0.0021 5.5489 ±0.2307 0.1499 ±0.0056 1533 ±12 1908 ±79 2345 ±8712 113 16 0.7359 ±0.0048 0.4103 ±0.0030 7.9015 ±0.2560 0.1397 ±0.0045 2216 ±16 2220 ±72 2223 ±7213 23 4 0.6957 ±0.0221 0.5405 ±0.0039 14.3612 ±0.8254 0.1927 ±0.0115 2786 ±20 2774 ±159 2765 ±16514 49 3 0.6108 ±0.0023 0.021 ±0.0002 0.1723 ±0.0096 0.0594 ±0.0032 134 ±1 161 ±9 583 ±3217 39 5 1.6465 ±0.0146 0.3424 ±0.0033 5.4493 ±0.6628 0.1154 ±0.0143 1898 ±18 1893 ±230 1887 ±23518 110 16 0.5712 ±0.0053 0.4109 ±0.0031 8.2348 ±0.2830 0.1453 ±0.0050 2219 ±17 2257 ±78 2292 ±7819 511 27 0.4994 ±0.0092 0.0215 ±0.0002 0.1544 ±0.0009 0.052 ±0.0003 137 ±1 146 ±1 287 ±220 90 15 0.4619 ±0.0020 0.4452 ±0.0034 10.2039 ±0.3503 0.1662 ±0.0056 2374 ±18 2453 ±84 2520 ±8422 269 13 0.2707 ±0.0017 0.0214 ±0.0006 0.1453 ±0.0073 0.0491 ±0.0014 137 ±4 138 ±7 155 ±424 42 7 0.6048 ±0.0059 0.4738 ±0.0037 10.983 ±0.6249 0.1681 ±0.0095 2500 ±20 2522 ±143 2539 ±144

224 J.-F. Shen et al. / Lithos 156–159 (2013) 218–229

40Ar/36Ar often exceeds 45,000, and that of 40Ar/4He is generally withinthe range of 0.16–0.25 (Feng et al., 2006; Hu et al., 1999; Simmons et al.,1987; Stuart et al., 1995). Thus, the He–Ar isotopic composition in thecrust and mantle is significantly different. Furthermore, it is usuallyconsidered that the He–Ar isotopic composition in the atmosphericsaturated water is the same as that in the surface atmosphere, with3He/4He=1Ra, 40Ar/36Ar=295.5, and 40Ar/4He approximately 0.001(Hu et al., 1999; Simmons et al., 1987; Stuart et al., 1995). Thus, thesignificant differences between the crust and mantle reservoirs allow3He/4He, 40Ar/36Ar and 40Ar/4He to be used as potential traces in iden-tifying the source characteristics of ore-forming fluids.

The pyrites from BMH iron deposit possess 3He/4He ratios of 0.0648–0.1886Ra, with a mean value of 0.1237Ra (Table 5), which are a slightlyhigher than that of the crust (0.01–0.1 R/Ra), but markedly lower thanthat of the mantle (6–9 R/Ra). On the 3He versus 4He diagram(Fig. 11), all of the data points from pyrites plot between the crust andmantle, but closer to crustal domain. The percentage of mantle-derivedHe can be calculated according to the crust–mantle mixing model,which is expressed as: He=[(3He/4He)(sample)−(3He/4He)(Crust)]/[(3He/4He)(Mantle)−(3He/4He)(Crust)]×100 (Xu et al., 1995).

Table 3Results of zircon LA-ICP-MS dating on sample B245-1.

Spot U(ppm)

Th(ppm)

232Th/238U

Error 206Pb/238U

Error 207Pb/235U

Error

1 59 91 1.5413 ±0.1380 0.3109 ±0.0074 4.5998 ±0.11222 33 64 1.9301 ±0.0187 0.3272 ±0.0029 5.0344 ±0.04793 251 244 0.9734 ±0.0049 0.4716 ±0.0034 10.6336 ±0.05424 491 265 0.5399 ±0.0075 0.4761 ±0.0034 10.8433 ±0.05735 230 71 0.3086 ±0.0046 0.4765 ±0.0032 10.8165 ±0.04886 135 156 1.1519 ±0.0107 0.4704 ±0.0034 10.377 ±0.05728 126 72 0.5688 ±0.0037 0.4927 ±0.0029 11.8569 ±0.041710 321 215 0.6686 ±0.0040 0.4735 ±0.0037 10.868 ±0.141111 6 5 0.814 ±0.0064 0.3439 ±0.0028 5.6294 ±0.190113 99 85 0.858 ±0.0044 0.4698 ±0.0034 10.4637 ±0.058514 196 101 0.5152 ±0.0047 0.5026 ±0.0038 12.2847 ±0.079515 88 92 1.0451 ±0.0078 0.4627 ±0.0034 10.1976 ±0.067216 19 49 2.5815 ±0.0407 0.4886 ±0.0032 11.8199 ±0.100117 257 299 1.1623 ±0.0216 0.4658 ±0.0035 10.3631 ±0.060118 266 142 0.5325 ±0.0075 0.3334 ±0.0049 7.3404 ±0.101122 158 58 0.3662 ±0.0177 0.2607 ±0.0076 4.9464 ±0.132823 635 294 0.4627 ±0.0042 0.0241 ±0.0003 ±0.2002 0.004224 734 140 0.191 ±0.0057 0.4653 ±0.0031 10.3334 ±0.051425 46 19 0.4188 ±0.0039 0.4608 ±0.0085 9.9622 ±0.170027 673 101 0.1496 ±0.0031 0.3876 ±0.0026 8.441 ±0.040128 450 95 0.211 ±0.0063 0.3331 ±0.0027 5.2205 ±0.031829 234 134 0.5742 ±0.0100 0.4726 ±0.0029 10.7701 ±0.040330 205 110 0.5373 ±0.0204 0.0431 ±0.0003 0.8334 ±0.0107

Where the lower limit of 3He/4He of the crust end-member is2×10−8 and that of the mantle is 1.1×10−8 (Stuart et al., 1995).The results show that the percentage of the mantle derived He inthe pyrite from the BMH deposit ranges from 0.17 to 2.98. In general,the 3He/4He ratios of ore fluids in the BMH are close to the crustalvalue, reflecting that the ore fluids mainly came from the crust andwere mixed with a small amount of the mantle component in themetallogenic process.

In addition, the 40Ar/36Ar ratios mostly range from 311.7 to 673.5,except for one sample which gives 22,909.4 (Table 5 and Fig. 12). Theresults are higher than that of the atmosphere (40Ar/36Ar=295.5),indicating the presence of excess argon produced probably by higherradiogenic 40Ar.

But as we know, the 40Ar/4He ratios of mantle fluids are in therange of 0.33–0.56 (Dunai and Touret, 1995) and the average valueof the crust is ~0.2 (Hu, 1997; Stuart et al., 1995). The 40Ar/4He ratiosof ore fluids in the BMH iron deposit are 0.362–0.716 (see Table 5),close to the value of the mantle.

However, the plots of 3He/4He (R/Ra) vs. 40Ar/36Ar of fluids in py-rite from the BMH iron deposit (Fig. 13) fall close to the field of the

207Pb/206Pb

Error 206Pb/238U(age)

Error 207Pb/235U(age)

Error 207Pb/206Pb(age)

Error

0.1073 ±0.0009 1745 ±41 1749 ±43 1754 ±150.1116 ±0.0009 1825 ±16 1825 ±17 1826 ±140.1635 ±0.0009 2491 ±18 2492 ±13 2493 ±140.1652 ±0.0009 2510 ±18 2510 ±13 2509 ±140.1646 ±0.0009 2512 ±17 2507 ±11 2504 ±140.16 ±0.0009 2485 ±18 2469 ±14 2456 ±130.1745 ±0.0010 2582 ±15 2593 ±9 2602 ±150.1665 ±0.0017 2499 ±19 2512 ±33 2522 ±260.1187 ±0.0039 1905 ±16 1921 ±65 1937 ±630.1615 ±0.0009 2483 ±18 2477 ±14 2472 ±140.1773 ±0.0010 2625 ±20 2626 ±17 2628 ±140.1598 ±0.0009 2452 ±18 2453 ±16 2454 ±140.1754 ±0.0018 2565 ±17 2590 ±22 2610 ±260.1614 ±0.0009 2465 ±19 2468 ±14 2470 ±130.1597 ±0.0009 1855 ±27 2154 ±30 2452 ±130.1376 ±0.0027 1494 ±44 1810 ±49 2197 ±430.0603 ±0.0007 153 ±2 185 ±4 613 ±80.1611 ±0.0009 2463 ±17 2465 ±12 2467 ±130.1568 ±0.0010 2443 ±45 2431 ±41 2421 ±160.158 ±0.0009 2111 ±14 2280 ±11 2434 ±130.1137 ±0.0006 1853 ±15 1856 ±11 1859 ±100.1653 ±0.0009 2495 ±15 2503 ±9 2510 ±140.1401 ±0.0017 272 ±2 615 ±8 2229 ±27

Fig. 8. U–Pb concordia diagram of zircons in weakly altered diorite porphyrite(a. B110-9-2) and in skarnization diorite (b. B245-1).

Fig. 9. The distribution of pyrite δ34S‰ in the Beiminghe iron ore deposit comparedwith those from other localities.

225J.-F. Shen et al. / Lithos 156–159 (2013) 218–229

crustal fluid component, indicating that the ore fluids were mainlyderived from crust.

In summary, it is inferred that the ore-forming fluids of BMH irondeposit are mainly derived from crust with small volume of mantlecomponent during the metallogenic processes.

5. Discussion

5.1. Timing of mineralization and magmatism in the Beiminghe irondeposit

In skarn-type ore deposits, the timing of mineralization is broadlycoincident with the time of magma emplacement. The 40Ar/39Ar ageof 137±2 Ma obtained from phlogopite in the skarn iron ores ofpresent study is markedly consistent with the U–Pb age of 136±2 Maobtained from zircons in the weakly altered diorite porphyrite

Table 4δ34S values of pyrite in Beiminghe Fe-ore deposit.

Sample no. Tested mineral δ34S‰

B-110-6-3 Pyrite 13.5B-230-2-5 Pyrite 12.2B-230-2-3 Pyrite 13.7B-245-3 Pyrite 14.4B-245-5 Pyrite 16.5

associated with the mineralization. Therefore, we conclude that timingof both magmatism and mineralization in the Beiminghe iron deposittook place at round of 136–137 Ma.

5.2. Sources of ore-forming materials

The sulfur isotope data obtained in our study indicate that theore-forming materials were mainly tapped from crustal sources by theMesozoic pluton. In the evolution models of 207Pb/204Pb with 206Pb/204Pb, the data fall in between the orogenic belt and mantle evolutionlines, mostly tending towards the orogenic belt line. The 208Pb/204Pband 206Pb/204Pb relationship shows a lower crustal origin. Thus, weinfer a mixed source, with most of the Pb derived from the lowercrust with limited input from mantle sources.

The 3He/4He and 40Ar/4He values indicate that the ore-formingfluids had a dominantly crustal origin, mixed with minor amounts ofmantle-derived fluids. The He–Ar isotopic composition, however pointsto an ultimate mantle origin for the fluids. However, in the evolved andhomogenized fluids system, the mantle-derived component is comput-ed to be less than 3%. The ore fluids preserve a complex evolutionaryhistory, including the involvement of minor meteoric water.

The summary of zircon U–Pb geochronology in our study indicatesthe following main age groups: (1) inherited Neoarchean andPaleoproterozoic magmatic zircons with ages ranging from 2500 to2786 and 2216 to 2408 Ma; (2) few Paleo-Mesoproterozoic metamor-phic zirconswith ages of 1553 Ma; and (3)magmatic zirconswith crys-tallization age of ca. 136 Ma. Obviously, the type (1) and (2) zirconsabove were captured from the country rocks during magma ascentand emplacement and confirms the varying degrees of incorporationof the Neoarchean and Paleo-Mesoproterozoic basement lithologiesduring magma tectonics. The Neoarchean and Paleoproterozoicinherited magmatic zircons in our samples from the Mesozoic oredeposit correlate well with similar ages widely reported from the base-ment rocks of the NCC (Wang and Liu, 2012;Wang et al., 2012; Zhai andSantosh, 2011). The Paleo-Mesoproterozoic metamorphic zirconscaptured by themagma are also in accordancewith similar aged zirconswidely reported from the high grade metamorphic rocks of the NCC.Thus, it is clear that the ore-forming magma had interacted with theancient lower crust of the NCC.

In summary the primary fluids appear to have been derived frommantle sources, which subsequently underwent a complex evolution-ary history. The high temperature magma, as it migrated throughthe lower crust, caused partial melting of the ancient lower crust,and older zircons were captured from the surrounding rocks. Whenthe hot magma was emplaced into the carbonate rocks, skarn

Fig. 10. Pb isotope composition of pyrite and host magmatic rock from the BMH iron deposit.

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mineralization occurred, generating the BMH iron deposit, as wellas similar occurrences in the Han-Xing region through contactmetamorphism.

5.3. Geodynamic implications

Santosh (2010) presented a synthesis of geological and geophysicalfeatures of the NCC including a re-interpretation of the deep seismicdata from various traverses across the crustal blocks and suture zones inthis Craton. The results brought out a multiple subduction history in thePaleoproterozoic with the Western and Eastern Blocks amalgamatedalong the Trans-North China Orogen (Central Orogenic Belt) and theYinshan and Ordos blocks assembled along the Inner Mongolia SutureZone, marking the final cratonization of the NCC. The sub-continental

Fig. 11. He isotope composition of BMH iron deposit.

mantle lithosphere beneath some of the old cratons of the world aremore than 200–250 km in thickness, and some of these have survivedfor periodsmore than 3 billion years in spite of later tectonic disturbance(Carlson, 2005; Gao et al., 2009; Pearson, 1999; Sleep, 2005; Zhu et al.,2011). The cratonic root of the NCC, since its formation in thePaleoproterozoic remained largely stable until the late Paleozoic (Xu etal., 2009a; Zhai and Bian, 2001; Zhao et al., 2001; Zheng and Wu, 2009),with an estimated thickness of over 200 km (Deng et al., 1994; Fan andMenzies, 1992). Activemagmatism in theEasternBlock of theNCC startedduring early Paleozoic triggering decratonization and refertilization,particularly beneath the Eastern Block of the NCC (Zhang, 2009, 2012;Zhang et al., 2012). The late Mesozoic witnessed extensive andcraton-wide magmatism resulting in the widespread destruction of thesub-continental lithospheric mantle, with a thin and fertile lithospherereplacing the older thick and refractory sub-continental mantle(e.g., Gao et al., 2002; Menzies et al., 1993; Santosh, 2010; Zhai et al.,2004). Thus, Mesozoic magmatism is a hallmark of craton destruction inthe NCC, and is therefore significant in understanding the geological his-tory and geodynamic processes associated with decratonization.

A number of studies have addressed the process of destruction ofthe NCC. Delamination (Gao et al., 2009) or erosion (Xu, 2006; Xuet al., 2009a) is among the principal mechanism suggested. Thermaland chemical erosion occurs through the upwelling of the hot as-thenosphere, which rises up and infiltrates the lithospheric mantleand lower crust (Zhang et al., 2005; Zhu et al., 2008). The heat andvolatile input would also lead to upper crustal remelting, therebygenerating magmas at various levels of the crust, with partial inputfrom the mantle (Luo et al., 1997). Extensive magma migrationfrom depth to shallower crustal levels has been considered as oneof the hallmarks of lithospheric destruction (Xu et al., 2009b).Subduction–erosion is also one of the important mechanisms bywhich extensive erosion of the ‘tectosphere’ (sub-continental mantlelithosphere) occurs (Santosh, 2010). Fluids released from downgoing

Table 5Helium and argon isotope data of pyrite from Beiminghe iron deposit.

Sample no. B110-6-3 B230-2-3 B230-2-5 B245-3 B245-5

Lithology Ore Skarn Ore Skarn Ore

4He (10−7) 2.37±0.16 4.97±0.34 10.87±0.73 3.87±0.26 15.30±1.03He/4He 0.1405 0.0907 0.2429 0.1280 0.26403He (10−7) 0.3327 0.4508 2.6403 0.4954 4.0392R/Ra 0.10034±0.00057 0.06479±0.00074 0.17350±0.0024 0.09140±0.0031 0.18860±0.001640Ar (10−7) 1.194±0.097 1.800±0.13 4.630±0.34 2.770±0.21 5.980±0.4440Ar/36Ar 359.2±13.7 673.5±041.4 22909.4±181.0 311.7±15.7 638.8±29.136Ar (10−7) 0.0033 0.0027 0.0002 0.0089 0.009440Ar⁎(10−7) 0.2188 1.0021 4.5709 0.1400 3.202340Ar⁎/4He 0.092 0.202 0.421 0.036 0.209F4He 4339.4 11122.3 328398.8 2627.4 9834.8

F4He=ratios of 4He/36Ar sample and 4He/36Ar atmosphere (4He/36Ar atmosphere=0.1655); 40Ar* is the excess argon without 40Ar atmosphere; 40Ar*=(40Ar)sample−295.5×(36Ar)sample.

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slabs hydrate and weaken the large mantle wedge, and magmasgenerated by slab melting rise up and invade the lithospheric mantle.The prolonged subduction of the Pacific slab from the east is consid-ered to have hydrated the mantle beneath East Asia and extensivelydestroyed the eastern part of the NCC.

The TM Belt records strong and widespreadmagmatic activity duringMesozoic, covering the whole “Yanshanian” event, with a peak at 124–153 Ma (Luo et al., 1996, 2006). Among these, the magmatic rocks ofsouthern TM were formed mainly during 125–136 Ma (Chen et al.,2005a; Dong et al., 2003; Peng et al., 2004; Wang et al., 2006; Zheng etal., 2007c). Their REE characteristics including lack of obvious Eu anoma-lies and the Sr–Nd–Pb isotopic features have been interpreted as the sig-nature of deep magma sources (Cai et al., 2004a, 2006; Niu et al., 1995).Furthermore, their low initial strontium ratio (w (87Sr)I/w (86Sr)I=0.7050–0.7069) has provided conclusive evidence for the involvementof mantle-derived materials mixing during formation in this area (Luoet al., 1997).

The time of formation of the BMH iron deposit was at ca. 136 Ma isconsistent with the peak of 125–136 Ma for the magmatic activity insouthern TM. Our study of the BMH iron ore deposit reveals mixedsource characteristics and complex evolution of the ore-formingfluids. The S, Pb, He and Ar isotopic tracers indicate that theore-forming materials were mainly sourced from the lower crust,with input from mantle-derived fluids. It is possible that the 136 Maage marks the initiation of the destruction of the NCC, as alsosupported the higher degree of crustal contamination. This featureis comparable with the processes associated with several majormetallogenic provinces of the world where mantle material transmit-ted to the crust involves multiple contaminations at various crustallevels (e.g., Mao et al., 2005; Zhao et al., 2000). Importantly, themetallogenic process of BMH iron deposit is also a significant markerfor the lithospheric destruction in this region.

Fig. 12. He–Ar isotope composition of BMH iron deposit.

The ore formation at BMH not only coincides with the time ofmagmatic activity in southern TMs, but also relates to the peak of110–140 Ma identified for the large-scale lithospheric destruction ofthe eastern block of NCC (Lin et al., 2008; Peng et al., 2004; Xu,2006; Zhu and Zheng, 2009; Zhu et al., 2011). The geodynamic settingfor magma generation and lithospheric destruction can be correlatedwith the tectonic processes associated with the deep and prolongedsubduction of the Pacific plate, and resulting tectonic, thermal andchemical erosion as well as refertilization through magma injection(Deng et al., 2000; Santosh, 2010; Tang et al., 2012; Xu et al., 2009a,2009b; Zhang, 2012).

A recent study on the zircons in granulite xenoliths entrained inCenozoic basalts from within the Trans-North China Orogen thatamalgamates the Western and Eastern Blocks of the NCC showed alarge spread of Phanerozoic concordant ages ranging from 470 Mato 40 Ma with peaks at 315 Ma, 220–230 Ma, 120 Ma and 46 Ma,suggesting episodic magmatic underplating in the ancient lowercrust of the NCC, lasting continuously throughout Phanerozoic, pro-ducing zircons from the underplated magmas or providing the heatsource for the recrystallization of zircons from the ancient crust (Liuet al., 2012; Zhang, 2012; Zhang et al., 2012). In another recentwork, Tang et al. (2006, 2012) investigated the high Mg peridotite xe-noliths in the Cenozoic Hebi basalts and the results indicate that thelithospheric mantle beneath the TNCO formed during the Archean andwas refertilized bymultiple additions of fluids andmelts. The critical lo-cation of the MBH iron ore deposit at the eastern margin of the TNCO,along the western periphery of the Eastern Block, and its genetic linkwith the Mesozoic magmatism in this region suggest that the mineral-ization is a response to the processes associated with lithosphericthinning in the NCC. It has been noted in some of the previous studiesthat the large-scale lithosphere thinning process in eastern China was

Fig. 13. 40Ar/4He vs. Rc/Ra plots for BMH iron deposit.

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accompanied by large-scale mineralization events (Yang et al., 2003;Zheng et al., 2007b), and the formation of BMH iron deposit is an obvi-ous manifestation of the NCC lithosphere thinning in southern TM.

6. Conclusion

1. Based on phlogopite 39Ar–40Ar dating and zircon U–Pb dating, it canbe determined that BMH iron deposit was formed in 136–137 Ma.

2. The ore-formation of BMH iron deposit mainly involved contribu-tions from the Late Archean and Early Proterozoic basement rocks.The inherited zircon ages and S, Pb, He–Ar isotope compositionsreveal that the fluids of BMH iron deposit were mainly derivedfrom the crust, with subordinate contribution from mantle fluids.We envisage lower crustal re-melting and pervasive crustal contam-ination through heat and fluid input frommantle source, thus yield-ing a mixed source signature for the metallogenic process.

3. Since the main ore-forming process of BMH iron deposit correlateswith the timing of destruction of the North China Craton, themetallogeny is considered as a manifestation of the destructionof the cratonic keel of the NCC below the southern TM region.

Acknowledgment

We thank Editor-in-Chief Prof. Nelson Eby and two referees fortheir constructive and helpful comments which greatly improvedthis paper. This study was financed by the National Natural ScienceFoundation of China (no. 90914002). We thank Hu Yuanyue, LiAiguo, and Zhang Dongbin (Minmetals Hanxing Mining Co., Ltd.) forproviding support in the field, access to the mines and data.

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