recycled crustal zircons from podiform chromitites in the luobusa ophiolite, southern tibet

Thematic ArticleRecycled crustal zircons from podiform chromitites in the

Luobusa ophiolite, southern Tibet

SHINJI YAMAMOTO,1* TSUYOSHI KOMIYA,1 HIROSHI YAMAMOTO,2 YOSHIYUKI KANEKO,3

MASARU TERABAYASHI,4 IKUO KATAYAMA,5 TSUYOSHI IIZUKA,6 SHIGENORI MARUYAMA,7

JINGSUI YANG,8 YOSHIAKI KON,9 AND TAKAFUMI HIRATA10

1Department of Earth and Astronomy Graduate School of Arts and Sciences, The University of Tokyo, Tokyo153-8902, Japan (email: [email protected]), 2Department of Earth and Environmental Sciences,Kagoshima University, Kagoshima 890-0065, Japan, 3Department of Education, Meisei University, Tokyo

191-8506, Japan, 4Department of Safety Systems Construction Engineering, Kagawa University, Takamatsu,Kagawa 761-0396, Japan, 5Department of Earth and Planetary Science Systems, Hiroshima University,

Kagami-yama 1-3-1, Higashi-Hiroshima, Hiroshima 739-8526, Japan, 6Department of Earth and PlanetaryScience, University of Tokyo, Hongo 7-3-1, Bunkyo, Tokyo 113-0033, Japan, 7Department of Earth and Planetary

Sciences, Tokyo Institute of Technology, 2-12-1 Ookayama, Meguro, Tokyo 152-8551, Japan, 8Key Laboratoryfor Continental Dynamics of the Ministry of Land and Resources, Tibet Center for Continental Dynamics,

Institute of Geology, Chinese Academy of Geological Sciences, Beijing 100037, China, 9Geological Surveyof Japan, National Institute of Advanced Industrial Science and Technology, 1-1-1, Higashi, Tsukuba,

Ibaraki 305-8567, Japan, and 10Laboratory for Planetary Sciences, Kyoto University, Kitashirakawa Oiwakecho,Kyoto 606-8502, Japan

Abstract We have measured the U–Pb age of zircon grains separated from podiformchromitites from the Luobusa ophiolite, Southern Tibet, using laser ablation microprobe –inductively coupled plasma mass spectrometer (LA-IC-PMS), to determine the age rela-tionship between the podiform chromitites and the host mantle peridotite. Spot analyseswith LA-IC-PMS, assisted by cathodoluminescence images gave a wide age range, fromthe Cretaceous to the Late Archean (ca 100–2700 Ma). The minimum ages of ca 100 Ma,plotted on the concordia curve, were slightly lower than the metasomatic (magmatic) eventin the supra-subduction zone (120 � 10 Ma), suggesting that the zircons suffered some Pbloss. However, most of the ages found are much older than those of the chromitite andophiolite formation. Laser Raman spectroscopy analyses revealed that the zircons recov-ered from the chromitites contain crustal mineral inclusions, such as quartz and K-feldspar,but lack mantle minerals (e.g., olivine, pyroxene, and chromite), suggesting that they hada crustal origin. The results indicate that crustal zircons in chromitites had a xenocrysticorigin and resided in the mantle peridotite for a long period before being entrained into thechromitite during its formation. This indicates that the mantle peridotite under the Neo-Tethys Ocean was affected by the crustal material contamination. Our results are consis-tent with previous reports that mid-oceanic ridge basalts in the Indian Ocean have theisotopic signature of crustal material contamination. From these results, and previousisotopic studies on Gondwana geology, we conclude that ancient zircons from podiformchromitites could provide evidence of crustal material being recycled through the uppermantle.

Key words: crustal contamination, Gondwana breakup, Neo-Tethys mantle, podiformchromitite, zircon U–Pb dating.

*Correspondence.

Received 1 March 2012; accepted for publication 7 October 2012.

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Island Arc (2013) 22, 89–103

© 2013 Wiley Publishing Asia Pty Ltd doi:10.1111/iar.12011

INTRODUCTION

Podiform chromitites generally occur in the upperpart of ophiolitic peridotites, especially inharzburgite-dominant ophiolites (e.g., Nicolas &Al Azri 1990). Processes such as magma mixingand melt–rock interaction play an important rolein most podiform chromitite genesis concepts(Lago et al. 1982; Zhou et al. 1996; Arai 1997;Matveev & Ballhaus 2002). Several of the studieson the genesis of podiform chromitites are basedon geochemical examination of chromitites andmineral inclusions within chromites (Augé 1987;McElduff & Stumpfl 1990; Zhou et al. 1996, 2005;Arai & Matsukage 1998; Bai et al. 2000; Malitch2004). It has long been recognized that podiformchromitites and their host peridotite can providesignificant information on melt–mantle reactionprocesses, but the origin and history of podiformchromitites are still controversial with tectonicsettings ranging between the supra-subductionzone and the mid-ocean ridges (Roberts 1988;Nicolas 1989; Arai & Matsukage 1998; Ballhaus1998). One of the reasons for these controversiesis the lack of precise chronological studies of podi-form chromitites.

Zircon is the mineral most widely used for theU–Pb age determination, and has proven to beuseful for dating both metamorphic and mantlerocks (e.g., Hoskin & Schaltegger 2003). Althoughzircon is rare in podiform chromitite, Grieco et al.(2001) reported zircons from podiform chromititesin phlogopite peridotite from the Finero complex,and their interpretation was that the zircon wasformed at the time of magmatism (metasomatism)responsible for the formation of the podiformchromitite. Recently, sensitive high-resolution ionmicroprobe (SHRIMP) U–Pb zircon ages havebeen reported from several studies of podiformchromitite in the Luobusa ophiolite, ranging from1657 to 457 Ma (Yang et al. 2001; Robinson et al.2009, 2011). However, the relationship between thevarious ages and the genesis of chromitite is notclear and is still controversial because the Luobusaophiolite, which includes the podiform chromitites,is thought to have formed in the Jurassic or EarlyCretaceous (Zhou et al. 1996, 2005; Robinson et al.2004; Shi et al. 2007; Yamamoto et al. 2009a). Thisuncertainty exists because minor systematicdating work has been performed on U–Pb zirconsfrom the Luobusa podiform chromitite.

We have focused on the systematic dating ofzircon U–Pb from podiform chromitite with spotlaser ablation microprobe – inductively coupled

plasma mass spectrometer (LA-ICP-MS) analysesof discrete domains within zircon crystals, assistedby cathodoluminescence (CL) images, because dif-ferent zircon domains may represent differentmetamorphic and tectonic histories. The abun-dance of zircon in podiform chromitite was low.Hence, we performed the heavy mineral separa-tion of five tons of podiform chromitite, and hand-picked more than 100 zircon grains. In addition,mineral inclusions in zircons from the podiformchromitites were carefully studied using laserRaman spectroscopy and an electron probemicroanalyzer (EPMA) to study their origin.

Here, we report the ages of zircons from podi-form chromitites in the Luobusa ophiolite rangingfrom the Cretaceous to the Late Archean (ca 2700–100 Ma), which contain crustal mineral inclusions(quartz, K-feldspar, apatite, and mica), suggestinga crustal origin for the zircons. Because of the agesrecorded for the zircons, combined with previousisotopic studies of the peridotite and basalt and therelationship with the regional geology, we discussthe origin of the zircons and the complex mantlehistory of the Neo-Tethys Ocean, including thehistory of the Luobusa ophiolite.

GEOLOGICAL OUTLINE OF THE LUOBUSA OPHIOLITE

The Luobusa ophiolite is located about 200 kmeast–southeast of Lhasa in southern Tibet (Fig. 1),in the Indus–Yarlung Zangbo suture zone (IYSZ),which separates the Indian subcontinent from theLhasa terrane (e.g., Allègre et al. 1984). It mainlyconsists of harzburgite and dunite, stratigraphi-cally overlain by a mélange that includes blocksof pillow lava, gabbro, pyroxenite, and chert in aserpentinite matrix. The northern edge of theLuobusa ophiolite is in fault contact with the Paleo-gene Luobusa Formation. To the south, Triassicflysch-type sedimentary rocks (Indian block)structurally overlie the ophiolite with a faultcontact. The northern boundary dips gently to thesouth, the southern boundary is a steep reversefault, and the ophiolite unit is stratigraphicallyoverturned (Yamamoto et al. 2007).

The mantle sequence mainly consists of spinel–harzburgite with or without clinopyroxene (Zhouet al. 1996; Xu et al. 2011). Harzburgite has a rela-tively high Mg# [= Mg/(Mg + Fe2+) ¥ 100] (90–92)and a low Al2O3 content (0.25–0.97 wt%). It showssystematic variations in CaO and Al2O3 contents,which are consistent with those of a residue of highlevels of partial melting (Zhou et al. 1996). These

90 S. Yamamoto et al.

© 2013 Wiley Publishing Asia Pty Ltd

characteristics are similar to some peridotites fromophiolite belts. The Cr# [= Cr/(Cr + Al) ¥ 100] ofthe accessory chrome spinel in harzburgite has awide range (22–67), is negatively correlated withMg#, and falls within the range of abyssal peridot-ites (Dick & Bullen 1984).

Geochronological data on the Luobusa ophioliteis not comprehensive, but the Xigaze ophiolites,located about 200 km west of Lhasa in the IYSZ,were formed in the Cretaceous (120 � 10 Ma)determined by measurements of radiolaria inchert (Ziabrev et al. 2003), whole rock U–Pb analy-ses (Göpel et al. 1984), and zircon SHRIMP analy-ses (Malpas et al. 2003) of gabbro, and havesupra-subduction zone affinity (Aitchison et al.2000; Dupuis et al. 2005). The Luobusa ophiolite,like most others in the IYSZ, has melt–rock inter-action and mineral chemistry features that areindicative of modification in a supra-subductionzone (Zhou et al. 1996, 2005; Dupuis et al. 2005).However, the Luobusa ophiolite was formed in themiddle Jurassic (177 � 31 Ma), determined usingSm–Nd isochron data for whole rocks and separateminerals from a gabbro dyke, and has TethyanOcean-type Nd and Pb isotopic characteristics(Zhou et al. 2002). Therefore, it is believed that theLuobusa ophiolite had a two-stage origin: (i) a mid-

ocean ridge at 177 � 31 Ma; and (ii) a supra-subduction zone at 120 � 10 Ma (Zhou et al. 1996,2005; Robinson et al. 2004).

Podiform chromitite ores crop out sporadicallyat the stratigraphic upper level in the harzburgiteunit (Fig. 1) as lens-shaped bodies with duniteenvelopes that are assumed to be produced by thereaction between the mafic melts and the host peri-dotites (Zhou et al. 1996, 2005). They are almostconcordant with the direction of deformation in thehost peridotites. According to the modal amountand texture of chromite and olivine, the chromi-tites are classified into massive, disseminated,nodular, dunitic orbicular, and banded types. Themassive-type chromitite is the most dominant inthe study area, and each type of chromitite com-monly transects others, indicating their multi-stage magmatic origin (Zhou et al. 1996, 2005;Yamamoto et al. 2009a). Chromites from the podi-form chromitites have relatively uniform Cr# andMg# values, in the range 70–83 and 61–78, respec-tively. The most common matrix mineral in thechromitite is olivine, which is very highly magne-sian (Mg# = 92–98), probably because of Mg–Fesubsolidus exchange. Clinopyroxene is the mostcommon mineral inclusion in chromites, followedby olivine, amphibole, chlorite, serpentine, sul-

Fig. 1 Geological map of the Luobusaophiolite, slightly modified from Zhouet al. (1996) by the results of our fieldsurvey in 2001 and 2006. IYSZ, Indus–Yarlung Zangbo suture zone.

Recycled zircons in podiform chromitites 91


fides, and platinum-group element alloys. Themassive- and nodular-type chromites have moremineral inclusions than the disseminated- andbanded-type chromites.

ANALYTICAL PROCEDURES

The Luobusa ophiolite chromitite samples werecollected directly from outcrops and brought to thelaboratory to avoid contamination. Preliminarymineral separation of 5 tons of chromitites wascarried out at the Zheng Zhou Institute, HenanProvince, China. Detailed procedures were thesame as reported by Robinson et al. (2004). Later,the heavy fractions were extracted with a mineralseparation system using conventional heavy liquidand magnetic separation techniques at TokyoInstitute of Technology. Finally, zircon grains werehand-picked at random under a stereoscope andthen mounted in 10 mm epoxy resin discs. Onefraction each from ore-containing Luobusa ophio-lite site, sites 1 (Fig. 1) and 11 (located at theeastern end of the ophiolite; see the locality mappublished by Xu et al. 2009), were selected forzircon dating.

Heavy mineral separation of a 2 kg chromititesample (site 11 LA426; massive- to disseminated-type chromitite) was performed in Tokyo Instituteof Technology laboratory to determine the abun-dance of zircon within a chromitite. Because of thelow abundance of zircon within the heavy fractions,we used an ultraviolet fluorescence system to iden-tify them under a microscope (Fig. 2), and three

zircon grains were found. The original locations ofthe zircons in the chromitite sample are, therefore,uncertain.

Mineral inclusions in the zircons were investi-gated at Tokyo Institute of Technology. LaserRaman microspectroscopy (JASCO NRS-2000C)was performed using the 514.5 nm line from anAr-ion laser at 20–50 mW. Chemical compositionsof the mineral inclusions in the zircons were ana-lyzed using an EPMA (JEOL JXA8800) with anaccelerating voltage of 15 kV, 12 nA beam current,and counting times of 20–40 s. The oxide ZAF cor-rection method was applied to the mineral compo-sition calculations. The internal zoning pattern ofthe zircon crystals was investigated using CL, andspot analyses on zircons using LA-ICP-MS wereperformed at the same laboratory (which had thenmoved to Kyoto University). The locationsfor the spot analyses on the zircon grains wereselected from the CL images and photomicro-graphs (using transmitted light) to avoid contami-nation of mineral inclusions and cracks within thezircon grains. The LA-ICP-MS instrumental con-ditions and measurement procedures were thesame as those used by Rino et al. (2004), using aThermoElemental VG PlasmaQuad 2 quadrupole-based ICP-MS equipped with an S-option interface(Hirata & Nesbitt 1995; Hirata 1997, 2000; Apinya& Hirata 2004; Iizuka & Hirata 2004; Hirata et al.2005). The laser ablation system was a MicroLasproduction (Gottingen, Germany) GeoLas 200CQ,which uses a Lambda Physik (Gottingen,Germany) COMPex 102 ArF excimer laser as a193 nm deep ultraviolet (DUV) light source. The

Fig. 2 Microphotographs (open polar) of heavy mineral separates from the podiform chromitite under (a) transmitted light, and (b) ultraviolet light.Zircons glow slightly orange under ultraviolet light.



LA-ICP-MS instrumental sensitivity was1.5 ¥ 104 cps/mg for both Pb and U for theNIST 610 standard reference material (SRM)using a 16 mm diameter pit ablated with a 5 Hzrepetition rate and a source pulse energy of140 mJ. All measurements were carried out usingpeak jump acquisition mode at the peaks 202Hg,204Pb (204Hg), 206Pb, 207Pb, 232Th, and 238U. NIST 610and 91500 zircon standards were used for normal-izing 207Pb/206Pb and 206Pb/238U. Isoplot v3.7 soft-ware (Ludwig 2008) was used for the 207Pb/235U–206Pb/238U concordia diagrams and the calcu-lation of the 207Pb/206Pb ages.

We analyzed the cores and rims of 60 zircongrains from the chromitite ore bodies from site 1(21 grains), site 11 (36 grains), and the 2 kg rocksample from LA426 (three grains) usingLA-ICP-MS to examine their age relationships(Table 1). We accepted the core data when theU–Pb ratios for the core and rim of a single grainwere within a 2s error range.

RESULTS

INTERNAL STRUCTURE USING CL IMAGES AND MINERALINCLUSIONS WITHIN ZIRCONS

Zircons from the podiform chromitites in theLuobusa had a wide range of sizes (30–200 mm) andwere colorless or slightly brownish. They hadvarious crystal shapes from prismatic to subhedraland round, and some grains showed metamictiza-tion by lattice radiation damage.

In some cases, zircons contained several mineralinclusions (Fig. 3). Laser Raman and EPMAanalyses revealed that they were quartz, feldspar,mica, apatite, biotite, hematite, and ilmenite. Nohigh-pressure minerals (e.g., coesite and diamond)or mantle minerals (e.g., olivine, pyroxene, andchromite) were found within the zircons. The mostabundant mineral inclusion is apatite (>50% oftotal inclusions), and the second one is quartz(about 30%). The distribution of the mineral inclu-sions within zircons does not correlate with thezircon morphology, or with different CL domainsor U–Pb ages, as shown below.

The CL analyses revealed that the internalstructure of the zircons contained various zoningtextures, which were classified into three maingroups: oscillatory, homogenous, and complex(Fig. 4). The zircon internal zoning pattern fromthe CL images is used to distinguish empiricallybetween igneous and metamorphic zircons (as in,e.g., Corfu et al. 2003). Oscillatory zonation is gen-

erally interpreted as growth zoning because ofcrystallization after a melt, and homogeneouszonation and overgrowth are interpreted as meta-morphic origin.

More than 50% of the zircons displays oscilla-tory zoning, but clear oscillatory zoning is rare.Most grains have partly recrystallized domains(e.g., Fig. 4b,d,f) and exhibit weak or ambiguousoscillatory zoning, suggesting a rehomogenizedtexture. Some grains exhibit inconsistent oscilla-tory zoning depending on their outer morphology(Fig. 4b,c,f). These features suggest that theywere affected by secondary thermal overprints,corrosion, or resorption. About 30% of the zirconsshows homogeneous or weak zonation. The homo-geneous textured grains are usually round, butsome grains have a prismatic morphology(Fig. 4h,l). In rare cases, sector zoning wasobserved (Fig. 4j). The homogeneous types oftenshow metamorphic overgrowth in the rim(Fig. 4h,i). These zoning features are well knownin metamorphic rocks (e.g., Katayama et al. 2001;Corfu et al. 2003). Approximately 20% of zirconsdisplays complex zoning, and these grains mostlyhave rough surfaces or an irregular morphology(Fig. 4m,n,q), suggesting multiple growth–dissolution–recrystallization processes. The greatvariety in internal structures, including oscillatory,homogeneous, and complex zoning, suggests thatthe zircons were derived from multiple sourcesand/or were affected by multiphase magmatic/metamorphic events.

U–Pb DATING OF ZIRCON WITH LA-ICP-MS

The LA-ICP-MS data for the zircons are listed inTable 1. The results are also shown as 207Pb/235U–206Pb/238U concordia diagrams in Figure 5, in whichthe different apparent ages of the cores and rimswithin a single grain are connected with a dashedline. In general, younger zircons (<500 Ma) show alarge error in their 207Pb/206Pb ages because oftheir lower Pb content and the uncertainty inher-ent in this Pb correction. In this situation, 206Pb/238U ages give higher precision and are morereliable than the 207Pb/206Pb age. Therefore, the206Pb/238U zircon age population was calculated andis shown as a histogram (Fig. 6).

The LA-ICP-MS spot analyses show thatzircons from the chromitites have a wide apparentage range, and were formed from the Cretaceousto the Late Archean (ca 100–2700 Ma) (Figs 5 and6). In the zircon ages histogram three age groupswere identified: (i) a broad and slight peak in the



Table 1 U-Pb LA-ICP-MS isotopic analytical data for zircons from podiform chromitites

Sample Domain Th/U 206Pb†/238U (2s) 207Pb†/235U (2s) 207Pb†/206Pb† (2s) Age (Ma) ‡% conc206Pb/238U (2s) 207Pb†/235U (2s)

1–2m 1 rim 0.55 0.25414 � 0.01739 4.90447 � 0.34512 0.13996 � 0.00231 1460 � 89 1803 � 58 81.01 core 0.38 0.30879 � 0.02112 5.32236 � 0.37453 0.12501 � 0.00206 1735 � 103 1872 � 58 92.62 core 0.63 0.03419 � 0.00234 0.24764 � 0.01743 0.05253 � 0.00087 217 � 15 225 � 14 96.53 rim 0.55 0.03914 � 0.00228 0.31850 � 0.01943 0.05902 � 0.00036 247 � 15 281 � 20 88.23 core 0.56 0.04716 � 0.00275 0.38923 � 0.02374 0.05985 � 0.00036 297 � 18 334 � 24 89.04 core 0.57 0.03535 � 0.00242 0.25185 � 0.01772 0.05167 � 0.00085 224 � 15 228 � 14 98.25 rim 0.01 0.05266 � 0.00307 0.37609 � 0.02294 0.05180 � 0.00031 331 � 20 324 � 23 102.15 core 0.41 0.06405 � 0.00438 0.48380 � 0.03404 0.05478 � 0.00090 400 � 26 401 � 23 99.96 core 0.70 0.03702 � 0.00253 0.31211 � 0.02196 0.06115 � 0.00101 234 � 16 276 � 17 85.07 core 0.89 0.29175 � 0.01996 5.07277 � 0.35697 0.12610 � 0.00208 1650 � 99 1832 � 58 90.18 core 0.53 0.02650 � 0.00181 0.20409 � 0.01436 0.05585 � 0.00092 169 � 11 189 � 12 89.49 core 1.35 0.03918 � 0.00268 0.41115 � 0.02893 0.07612 � 0.00126 248 � 17 350 � 21 70.8

12 core 0.57 0.27728 � 0.01897 4.46513 � 0.31421 0.11679 � 0.00193 1578 � 95 1725 � 57 91.513 core 0.78 0.32680 � 0.02236 5.89531 � 0.41485 0.13083 � 0.00216 1823 � 108 1961 � 59 93.014 core 0.36 0.08653 � 0.00592 0.77457 � 0.05451 0.06492 � 0.00107 535 � 35 582 � 31 91.915 core 0.31 0.07046 � 0.00482 0.81572 � 0.05740 0.08396 � 0.00139 439 � 29 606 � 32 72.516 rim 0.01 0.03978 � 0.00255 0.26988 � 0.01747 0.04920 � 0.00016 251 � 16 243 � 18 103.716 mantle 2.11 0.06533 � 0.00418 0.47550 � 0.03078 0.05279 � 0.00017 408 � 27 395 � 31 103.316 core 2.02 0.06529 � 0.00418 0.50401 � 0.03263 0.05599 � 0.00018 408 � 27 414 � 33 98.417 core 0.44 0.31177 � 0.02133 5.25614 � 0.36987 0.12227 � 0.00202 1749 � 104 1862 � 58 94.019 core 0.20 0.06716 � 0.00459 0.60756 � 0.04275 0.06562 � 0.00108 419 � 28 482 � 27 86.922 core 0.74 0.03143 � 0.00183 0.24250 � 0.01479 0.05595 � 0.00034 200 � 12 220 � 15 90.523 rim 0.29 0.06609 � 0.00423 0.43502 � 0.02816 0.04774 � 0.00015 416 � 28 416 � 33 100.023 core 0.57 0.11570 � 0.00741 1.10037 � 0.07123 0.06898 � 0.00022 706 � 48 754 � 71 93.728 rim 0.21 0.28462 � 0.01947 4.69354 � 0.33028 0.11960 � 0.00197 1615 � 97 1766 � 57 91.428 core 0.39 0.33932 � 0.00724 5.28317 � 0.11400 0.11292 � 0.00036 1883 � 47 1866 � 110 100.929 core 0.84 0.05190 � 0.00355 0.41541 � 0.02923 0.05805 � 0.00096 326 � 22 353 � 21 92.530 core 0.62 0.03702 � 0.00237 0.28017 � 0.01814 0.05488 � 0.00018 234 � 15 251 � 18 93.4

11-2m 1 rim 0.32 0.22849 � 0.01202 4.96732 � 0.29518 0.15767 � 0.00145 1323 � 77 1799 � 281 73.51 core 1.02 0.41777 � 0.00732 9.53209 � 0.18881 0.16548 � 0.00153 2250 � 47 2391 � 176 94.12 core 0.17 0.07428 � 0.00391 0.58233 � 0.03460 0.05686 � 0.00052 462 � 25 466 � 35 99.14 core 0.51 0.11306 � 0.00595 1.00379 � 0.05965 0.06439 � 0.00059 690 � 38 706 � 60 97.85 rim 0.05 0.38783 � 0.00680 11.26397 � 0.22312 0.21065 � 0.00194 2113 � 44 2545 � 204 83.05 core 0.34 0.51013 � 0.00894 16.42664 � 0.32538 0.23354 � 0.00215 2657 � 57 2902 � 286 91.66 rim 0.71 0.06993 � 0.00368 0.57248 � 0.03402 0.05938 � 0.00055 436 � 24 460 � 34 94.86 core 0.83 0.14557 � 0.00766 1.43552 � 0.08530 0.07152 � 0.00066 876 � 49 904 � 85 96.97 core 0.67 0.02232 � 0.00105 0.15460 � 0.00781 0.05024 � 0.00031 142 � 7 146 � 8 97.58 core 0.78 0.07501 � 0.00353 0.68405 � 0.03454 0.06614 � 0.00040 466 � 23 529 � 35 88.19 core 0.91 0.28696 � 0.01350 6.22622 � 0.31442 0.15736 � 0.00096 1626 � 87 2008 � 304 81.0

10 rim 0.71 0.08277 � 0.00389 0.75964 � 0.03836 0.06657 � 0.00041 513 � 25 574 � 39 89.310 core 0.86 0.07614 � 0.00358 0.60966 � 0.03079 0.05808 � 0.00035 473 � 23 483 � 31 97.913 core 1.20 0.07346 � 0.00215 0.61232 � 0.02371 0.06045 � 0.00051 457 � 14 485 � 24 94.215 core 0.55 0.35323 � 0.01032 5.88803 � 0.22803 0.12090 � 0.00102 1950 � 66 1959 � 223 99.516 core 0.20 0.31478 � 0.00920 4.96193 � 0.19217 0.11432 � 0.00097 1764 � 59 1813 � 189 97.317 core 0.82 0.06882 � 0.00201 0.57298 � 0.02219 0.06039 � 0.00051 429 � 13 460 � 22 93.318 rim 0.47 0.07454 � 0.00318 0.56895 � 0.02571 0.05535 � 0.00028 464 � 20 457 � 26 101.418 core 0.18 0.10749 � 0.00459 1.46797 � 0.06633 0.09905 � 0.00049 658 � 29 917 � 67 71.722 core 0.55 0.04166 � 0.00177 0.33507 � 0.01591 0.05834 � 0.00041 263 � 11 293 � 16 89.723 core 0.53 0.08301 � 0.00353 0.65187 � 0.03094 0.05695 � 0.00040 514 � 23 510 � 31 100.925 core 2.37 0.01538 � 0.00065 0.10792 � 0.00512 0.05089 � 0.00036 98 � 4 104 � 5 94.628 core 1.53 0.19874 � 0.00845 2.31706 � 0.10999 0.08456 � 0.00059 1169 � 54 1218 � 110 96.0

11-2s 1 rim 0.05 0.11050 � 0.00471 0.99238 � 0.04485 0.06513 � 0.00032 676 � 30 700 � 45 96.51 mantle 0.08 0.15961 � 0.01140 1.55345 � 0.11410 0.07059 � 0.00040 955 � 73 952 � 114 100.31 core 0.12 0.16724 � 0.01195 1.64128 � 0.12056 0.07118 � 0.00041 997 � 77 986 � 120 101.12 core 0.29 0.07530 � 0.00538 0.62872 � 0.04618 0.06055 � 0.00035 468 � 35 495 � 47 94.53 core 0.31 0.07645 � 0.00546 0.57721 � 0.04240 0.05476 � 0.00031 475 � 35 463 � 43 102.65 rim 0.12 0.30129 � 0.01284 4.36612 � 0.19731 0.10510 � 0.00052 1698 � 83 1706 � 194 99.55 core 0.21 0.33302 � 0.00793 4.97622 � 0.12184 0.10838 � 0.00062 1853 � 51 1815 � 117 102.16 rim 0.14 0.07014 � 0.00300 0.53453 � 0.02415 0.05528 � 0.00027 437 � 19 435 � 24 100.56 mantle 0.21 0.08073 � 0.00577 0.70711 � 0.05194 0.06353 � 0.00036 500 � 37 543 � 52 92.26 core 0.34 0.16226 � 0.00693 1.96715 � 0.08889 0.08793 � 0.00044 969 � 45 1104 � 89 87.87 core 0.13 0.01899 � 0.00136 0.12616 � 0.00927 0.04819 � 0.00027 121 � 9 121 � 9 100.58 core 0.24 0.07369 � 0.00526 0.54734 � 0.04020 0.05387 � 0.00031 458 � 34 443 � 41 103.4

10 core 0.72 0.12203 � 0.00577 1.13347 � 0.05517 0.06737 � 0.00026 742 � 37 769 � 56 96.512 core 0.22 0.07800 � 0.00369 0.58810 � 0.02862 0.05469 � 0.00021 484 � 24 470 � 29 103.113 core 0.22 0.06866 � 0.00294 0.51747 � 0.02340 0.05466 � 0.00027 428 � 19 424 � 24 101.115 core 0.74 0.07617 � 0.00308 0.58563 � 0.02461 0.05576 � 0.00021 473 � 20 468 � 25 101.116 rim 0.07 0.07426 � 0.00300 0.58579 � 0.02461 0.05721 � 0.00022 462 � 19 468 � 25 98.616 core 0.25 0.11897 � 0.00481 1.00870 � 0.04238 0.06149 � 0.00024 725 � 31 708 � 43 102.317 core 0.48 0.01883 � 0.00076 0.11837 � 0.00497 0.03914 � 0.00015 120 � 5 114 � 5 105.918 core 0.43 0.14355 � 0.00580 1.36204 � 0.05723 0.06881 � 0.00026 865 � 37 873 � 58 99.120 rim 0.00 0.01879 � 0.00076 0.10611 � 0.00446 0.04096 � 0.00016 120 � 5 102 � 5 117.220 core 0.92 0.08302 � 0.00335 0.70892 � 0.02979 0.06193 � 0.00024 514 � 22 544 � 30 94.521 core 0.55 0.37097 � 0.01499 6.00441 � 0.25228 0.11739 � 0.00045 2034 � 96 1976 � 246 102.922 core 1.95 0.13504 � 0.00546 1.19969 � 0.05041 0.06443 � 0.00025 817 � 35 800 � 51 102.024 core 1.37 0.01606 � 0.00164 0.11075 � 0.01131 0.05000 � 0.00019 103 � 11 107 � 11 96.3

LA426 1 core 1.08 0.08747 � 0.00201 0.69753 � 0.02695 0.05784 � 0.00180 541 � 12 537 � 16 100.62 core 0.35 0.01582 � 0.00092 0.11078 � 0.00676 0.05079 � 0.00031 101 � 6 107 � 7 94.83 core 1.05 0.01507 � 0.00088 0.10194 � 0.00622 0.04907 � 0.00030 96 � 6 99 � 6 97.8

All errors are quoted at 2s level. † Common Pb corrected using 204Pb. ‡ % conc is percentage discordance defined as (206Pb/238U age)/(207Pb/235U age) ¥ 100.1–2m, zircons from site 1 chromitite body; 11-2m and 11-2s, medium and small zircon grains recovered from site 11 chromitite body, respectively; LA426, chromitite rock

sample from site 11 and recovered three grains in our laboratory.



Middle Proterozoic (ca 1500–2000 Ma); (ii) a sig-nificant peak around the Ordovician (ca 430–500 Ma); and (iii) a second high and broad peakfrom the Cretaceous to the Permian (ca 100–290 Ma). There is no obvious correlation betweenage and zircon shapes or mineral inclusions. Theoldest two zircons are highly discordant (Fig. 5a),and have different apparent ages for their coresand rims, indicating that these grains suffered Pbloss after formation. The grains formed around200–2000 Ma are situated both on and off the con-cordia curve, and the youngest grains (100–110 Ma) fit well on the concordia curve (Fig. 5c).Because of the wide distribution of zircon agesshown in the diagram, a single discordia line modelcan not be applied. The reason of the wide agerange may be that the zircons suffered one or morePb loss events and/or were derived from differentsources, including several age groups. The Th/Uratios, which range from 0.01 to 2.37, do not appearto correlate with the zircon CL images, morpholo-gies, or ages.

Multiple spot analyses within a single zircongrain generally yield the same age (Fig. 4), butsome grains have younger apparent ages for therim than for the core (Fig. 7). These zircons exhibitan overgrowth texture, with a core of bright lumi-nescence, surrounded by a relatively low lumines-cent rim in the CL images. However, a few grainsshow the opposite pattern with a low luminescentcore surrounded by a bright rim. The inner bound-ary of the core and rim domains is usually irregu-lar, with relatively small cores. The cores and rimsshowing different ages and different luminescencehave different Th/U ratios within a single grain(Table 1). Compositional differences suggestthat the cores and rims crystallized in different

chemical environments. Therefore, the old coresare interpreted as inherited domains.

To investigate the age relationships between thecores and rims, 26 data points were obtained from13 zircon grains that had inherited cores, and their206Pb/238U age population is shown in Figure 8. Theages of the cores are widely distributed from theTriassic (ca 230 Ma) to the Late Archean (ca2700 Ma), and the rim ages exhibit a similar varia-tion but with an intense peak around 420–480 Ma.For example, two zircon grains with cores of ages725 and 969 Ma have younger rims dated at 462and 437 Ma, respectively (Fig. 7b,c).

DISCUSSION

AGE RELATIONSHIPS

As described above, geochronological data on theLuobusa ophiolite and podiform chromitite are notcomprehensive. So far, it has been proposed thatthe Luobusa ophiolite originated from a mid-oceanridge spreading center at 177 � 31 Ma, and wasmodified by supra-subduction zone magmatism(metasomatism) at 120 � 10 Ma (Zhou et al. 1996,2002, 2005; Robinson et al. 2004; Dupuis et al.2005; Xu et al. 2011). In this study, we obtained awide range of zircon ages, from ca 100 to 2700 Ma,but the great variety of internal structures shownby the CL images suggests that the zircons werederived from multiple sources and were affectedby multiphase magmatic/metamorphic events.

If zircons derived from multiple sources wereaffected by multiple thermal events, the minimumage would be significant to constrain the lastthermal/magmatic event. We found the minimum

Fig. 3 Microphotographs (open polar) of zircons from the podiform chromitite and mineral inclusions that are apatite, quartz, K-feldspar, biotite,ilmenite, and chlorite identified using laser Raman spectroscopy.



Fig. 4 CL images of zircon grainsfrom the podiform chromitite. Analyticalspots (dashed circles) about 16 mm indiameter by LA-ICP-MS and 206Pb/238Uages are shown on the CL images. (a–f)Oscillatory zoning (including partlyrecrystallized and rehomogenizedzoning), (g–l) homogeneous zoning, and(m–r) complex zoning. Zircon grainsample numbers correspond to those inTable 1. Scale bars, 50 mm.



age to be 96 � 6 Ma, which fits on the concordiacurve. Young zircon grains of about 100–110 Mashow (i) prismatic or round crystal shapes; (ii)vague or complex zoning textures (e.g.,Fig. 4j,l,p,r); and (iii) no mineral inclusions (exceptapatite inclusion; Fig. 4r). These features aresimilar to those of zircons from podiform chromi-tites in the Finero phlogopite–peridotite (Griecoet al. 2001) and in the Voikar–Syninsky ophiolite(Savelieva et al. 2006, 2007), which were affectedby subduction-related magmatism (metasoma-tism). Therefore, we infer that the younger zirconsgrew from a melt (or fluid) responsible for thesupra-subduction zone magmatism (metasoma-tism) at ca 120 Ma. The zircons appear to beslightly younger than the timing of the supra-subduction zone event, suggesting that they havesuffered some Pb loss.

However, the interpretation of ages older than100 Ma is less straightforward. In general, thepodiform chromitites are thought to have formedfrom the reaction between melt and wallrock (peri-dotite) under a ridge/spreading center in a mid-ocean or island-arc setting (Nicolas & Al Azri 1990;Zhou et al. 1996, 2005; Arai 1997; Arai & Matsuk-age 1998; Ballhaus 1998; Matveev & Ballhaus2002). Recent osmium isotopic studies on theRu–Os–Ir alloys from Luobusa podiform chromi-tites indicate model age ranges from 197 to 270 Maand an intense peak at ca 235 Ma (Shi et al. 2007).These ages are consistent with the opening of theNeo-Tethys Ocean (Li & Powell 2001). Therefore,it is plausible that the Luobusa podiform chromi-tites originated in a mid-oceanic ridge setting.

If all zircons were formed from a magma or fluidduring supra-subduction zone magmatism (meta-somatism) or from magmatic processes under a

Fig. 5 Concordia diagrams for LA-ICP-MS U–Pb analyses of zirconsfrom the podiform chromitite in the Luobusa ophiolite. (a) From 0 to3000 Ma, (b) from 0 to 1000 Ma, and (c) from 0 to 600 Ma. Results areindicated with 2s errors. When the ages of the core and rim of a singlegrain were coincident within 2s errors, we adopted the core data. Differentapparent ages of the core and rim within a single grain are connected witha dashed line.

Fig. 6 Population histogram of podiform chromitite zircon 206Pb/238Uages. Width of each column is 80 Ma.



mid-oceanic ridge, their ages would be expectedto be from the Triassic to the Cretaceous (ca 120–230 Ma). However, many zircons from the Luobusapodiform chromitite have both Paleozoic and Pre-cambrian ages, and they often contain crustal min-erals such as quartz and K-feldspar. It is hardlypossible that the anomalously old crustal zirconscould have crystallized in the mantle section belowa mid-oceanic ridge or supra-subduction zone.These zircons are, therefore, interpreted as xenoc-rysts that were entrained into the podiformchromitites during their formation.

CLOSURE TEMPERATURE OF THE ZIRCON

The zircon age discordance clearly indicates Pbloss, but a surprising feature is the concordanceof some zircons, which presumably resided inthe uppermost mantle between the time of theircrystallization and their entrainment by the podi-form chromitite. Previous studies of the zirconU–Th–Pb system (Mezger & Krogstad 1997) sug-gested that Pb diffusion in the pristine latticeoccurs down to temperatures of at least 1000°C.Experimental work has shown that the closuretemperature of the U–Th–Pb system in naturalzircon is >900°C (e.g. Lee et al. 1997). Lee et al.1997 determined Pb diffusion coefficients at 900,1000, and 1100°C, and consequently concluded that100 mm zircon can retain its ‘memory’ at 900°C forat most 150 Ma.

However, many zircons that have retained con-cordant ages for a long time, even in high tem-perature environments, have been reported fromthe mantle and UHT metamorphic rocks. Forexample, Möller et al. (2002) showed that zirconcores preserved their pre-UHT growth ages andremained concordant despite high temperaturesduring metamorphism (>950°C) in the UHTgranulites. Moreover, anomalously old zirconshave been reported repeatedly from mantle rocks,suggesting that a specific chemical environment inthe mantle allows zircons to retain their U–Pbages (Kinny & Meyer 1994; Pilot et al. 1998; Pel-tonen et al. 2003; Schwartz et al. 2005; Polat et al.2009). The evidence suggests that the zirconU–Pb system can remain closed even atextremely high temperatures, and that the zircon

Fig. 7 CL images of zircon grains thathave an inherited old core and a youngerrim. The circles on the CL images areanalytical spots obtained using LA-ICP-MS. Data are indicated as 206Pb/238U ageswith 2s errors. (a–f) Zircon grain samplenumbers correspond to those in Table 1.

Fig. 8 Zircon population histogram of 206Pb/238U ages, which have aninherited old core and a younger rim. (a) Ages of rims, and (b) ages ofcores. Width of each column is 80 Ma.



closure temperature is probably higher than1000°C.

Kresten et al. (1975), however, have reportedthat some kimberlite zircons have 5% smaller unit-cell volumes than granitic zircons, and suggestedthat the contraction of zircon suppresses Pb diffu-sion and raises their closure temperature. In fact,Yu et al. (2001) have examined the unit-cell volumeof zircons from Luobusa podiform chromitites andreported that they have shorter interatomic Zr–Oand Si–O bond distances than granitic zircons. Assome zircons show age concordance, we alsosuggest that the zircon U–Pb closure temperatureis much higher under specific conditions than hasbeen estimated from some experiments, and thatthe older zircon population was able to retainradiogenic Pb over a long period even under themantle conditions.

ORIGIN OF THE CRUSTAL ZIRCONS ANDTECTONIC IMPLICATIONS

The inheritance of unusual old crustal zirconsin the mantle suggests that the mantle peridotite

under the Neo-Tethys Ocean contained crustalmaterials. Crustal material contamination in theupper mantle has been reported repeatedly instudies of mid-oceanic ridge basalts (MORBs),especially in the Indian Ocean (e.g., Hanan et al.2004). The geochemical characteristics of basaltsin the Tethyan ophiolite in the IYSZ are similar tothose of the present Indian MORBs (e.g., Mahoneyet al. 1998), and gabbro dykes in the Luobusaophiolite have Indian Ocean-type Nd and Pb iso-topic characteristics (Zhou et al. 2002). It has beenwidely recognized that MORBs in the IndianOcean have different isotopic compositions fromthose in the Atlantic and Pacific Oceans (e.g.,Dupré & Allègre 1983); alternative suggestions forthis difference are that the radiogenic signaturemight be the result of oceanic crust and sedimentcontamination in the upper mantle by (i) ancientsubduction processes (e.g., Rehkämper &Hofmann 1997), or (ii) delamination and detach-ment of continental material into the shallowmantle during rifting and breakup of Gondwana(e.g., Arndt & Goldstein 1989). Figure 9 shows aschematic illustration on the tectonic evolution of

Fig. 9 Schematic diagrams showingthe tectonic evolution of the Luobusaophiolite and the Neo-Tethys Ocean(modified from Allègre et al. 1984;Maruyama et al. 1989; Li & Powell 2001;Ziabrev et al. 2003; Veevers 2005). (a,b)From 650 Ma to 500 Ma Gondwanalandconsolidated and crustal materials werebrought into and mixed with the uppermantle. The Lhasa block and Indiancontinent were in the northern part ofGondwanaland, and at ca 500 Ma Gond-wanaland started to break up. (c) Fromthe Late Triassic to the Early Jurassic (ca200–230 Ma), the Neo-Tethys Oceansplit the Lhasa block from the Indiancontinent. (d) At ca 170 Ma the Luobusaophiolite was generated in the Neo-Tethys Ocean overlying the ‘polluted’mantle. (e) At ca 120 Ma the Luobusaophiolite was affected by secondarymagmatism (metasomatism) in a supra-subduction zone setting. (f) Finally, theLuobusa ophiolite was obducted onto theedge of the Indian continent, now occur-ring in the suture zone.



the Luobusa ophiolite and the Neo-Tethys Oceanbased on the paleogeographic reconstructions pre-viously proposed (Allègre et al. 1984; Maruyamaet al. 1989; Li & Powell 2001; Ziabrev et al. 2003;Veevers 2005). The Lhasa block and the Indiancontinent were once in the northern part of Gond-wanaland, and after the breakup of Gondwana-land, the Neo-Tethys Ocean was formed betweenthe Lhasa block and the Indian continent(Maruyama et al. 1989; Li & Powell 2001; Veevers2005). At around 170 Ma, the Luobusa ophiolitewas formed in the Neo-Tethys Ocean, overlyingthe ‘polluted’ mantle (Zhou et al. 2002). It is highlyprobable that around the same time the podiformchromitites were formed and assimilated thesecrustal zircons from the polluted upper mantleduring partial melting and melt–mantle reactionprocess under a mid-oceanic ridge. Subsequently,the oceanic lithosphere, including the podiformchromitites, could have been affected by secondarymagmatism (metasomatism) in a supra-subductionzone setting (Aitchison et al. 2000; Ziabrev et al.2003; Zhou et al. 2005).

According to this scenario, we can not obtaintruly discordant grain ages from the concordiacurve, because the recycled zircons have severaldifferent origins and histories. However, the highpopulation of the age distribution around 500 Main Figures 6 and 8 suggests that at least onethermal/tectonic event was related to the crystal-lization, Pb loss, or metamorphic overprint on thezircons. The time around 500 Ma coincides withthe end of the Gondwanaland consolidation(Veevers 2003). In nearby regions of the Indiancontinent, it was a time of collision of the Archean–Proterozoic crustal blocks in convergent marginsettings (Veevers 2003). It is widely thought thatmid-Proterozoic supracrustal rocks were sub-ducted and metamorphosed with the sinking ofthe oceanic lithosphere to mantle depths duringMiddle Cambrian times and that large volumes ofgranitoids intruded into the convergent blocks atthe same time (Yoshida et al. 2003; Veevers 2005;Yoshida & Upreti 2006). Therefore, the majorzircon age population dated around 500 Ma maycorrespond to an igneous and metamorphic eventaround East Gondwana in the Late Cambrian tothe Early Ordovician.

The wide variation in ages plotted on and off theconcordia may be due to continuous and concor-dant Pb loss, a multiple source origin, and a multi-phase secondary overprint. Our interpretation ofthe zircon ages supports models such as the mantlebeing polluted by crustal materials (continental

crust or sediment) (e.g., Rehkämper & Hofmann1997; Hanan et al. 2004; Scholl & von Huene 2007;Polat et al. 2009; Yamamoto et al. 2009b), andsecondary magmatism (metasomatism) in asupra-subduction zone setting (Malpas et al. 2003;Dupuis et al. 2005; Zhou et al. 2005) (Fig. 9).Therefore, we propose that ancient zircons fromthe podiform chromitites in the Luobusa ophioliteare evidence of crustal materials recycled throughthe upper mantle.

On the other hand, many unusual mineral sepa-rates have been reported from recent studies ofpodiform chromitite in the Luobusa ophiolite,including UHP minerals (microdiamond, coesite,and moissanite) and highly reduced minerals(native elements, alloys, carbides, silicides, andnitrides) (Bai et al. 1993, 2000; Robinson et al.2004; Yang et al. 2007; Dobrzhinetskaya et al. 2009;Trumbull et al. 2009; Xu et al. 2009). The UHPmineral separates were initially assumed to bexenocrysts captured by chromites that were crys-tallized from boninitic magma at shallow mantledepths (Bai et al. 1993, 2000; Robinson et al. 2004).Yamamoto et al. (2009a) reported unusual coesiteand clinopyroxene exsolution lamellae withinLuobusa podiform chromitite chromites. Theexistence of coesite exsolution lamellae withinchromite clearly indicate UHP evidence ofLuobusa chromitite itself having a deep-mantleorigin, whereas low-pressure type (normal)chromites without lamellae were simultaneouslyobserved in the Luobusa podiform chromitite.Therefore, Yamamoto et al. (2009a) concluded thatthe podiform chromitites at Luobusa retain evi-dence of their multi-stage development from aUHP environment to low-pressure magmatic pro-cesses. Low-pressure magma interaction with thehost peridotite can enhance the precipitation ofnew minerals (chromite and zircon) and scavengethe recycled older zircons, which resided in themantle peridotite. This melt–mantle reactionprocess is possible in both mid-oceanic ridge andsupra-subduction zone settings.

In conclusion, unusual old crustal zircons in thepodiform chromitite are xenocrysts that origi-nated from continental materials that werebrought into the upper mantle by delaminationand detachment during the breakup of Gondwana-land and/or by ancient subduction processes, suchas sediment subduction and subduction erosion.They mixed with the upper mantle, resided inmantle peridotite for a long period, and finallywere entrained from the peridotite into the podi-form chromitite during its formation by melt–rock



interaction. The ages obtained from LA-ICP-MSsuggest three discrete stages of zircon growth: (i)Proterozoic to Late Archean, corresponding to theage of supracrustal materials in Gondwanaland;(ii) Late Cambrian to Early Ordovician (ca450–500 Ma), corresponding to a second stagemetamorphic overprint and/or a new igneous crys-tallization event in Gondwanaland; and (iii) a finalevent corresponding to supra-subduction zonemagmatism (metasomatism) at ca 120 Ma.

The upper mantle peridotites, during theirhistory from Gondwanaland to the Neo-TethysOcean, clearly experienced a long and complex tec-tonic and compositional evolution. Fortunately, theearly burial and metamorphic histories are wellpreserved as inclusion assemblages within thezircons. As seen above, zircons can record complexevolutionary histories of crustal conditions pre-served even in mantle material. The discovery ofunusual old crustal zircons from podiform chromi-tites provides us with a great deal of informationabout the generation of chromitites and thedynamics of mantle processes.

ACKNOWLEDGEMENTS

We thank S Rino (Tokyo Institute of Technology)for LA-ICP-MS analytical help. The authorsappreciate the work of X Xu and R Shi during fieldwork in Tibet. Comments by B F Windley and P FRobinson improved the paper. S Y is supported bythe Japan Society for the Promotion of Science.

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recycled crustal zircons from podiform chromitites in the luobusa ophiolite, southern tibet

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