Mapping lithospheric boundaries using Os isotopes of mantle xenoliths: An example from the North China Craton Jingao Liu a,⇑ , Roberta L. Rudnick a , Richard J. Walker a , Shan Gao b , Fu-yuan Wu c , Philip M. Piccoli a , Honglin Yuan d , Wen-liang Xu e , Yi-Gang Xu f a Department of Geology, University of Maryland, College Park, MD 20742, USA b State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan 430074, China c State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, P.O. Box 9825, Beijing 100029, China d State Key Laboratory of Continental Dynamics, Department of Geology, Northwest University, Xi’an 710069, China e College of Earth Sciences, Jilin University, Changchun 130061, China f Key Laboratory of Isotope Geochronology and Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou 510640, China Received 28 December 2010; accepted in revised form 19 April 2011; available 27 April 2011 Abstract The petrology, mineral compositions, whole rock major/trace element concentrations, including highly siderophile ele- ments, and Re–Os isotopes of 99 peridotite xenoliths from the central North China Craton were determined in order to con- strain the structure and evolution of the deep lithosphere. Samples from seven Early Cretaceous–Tertiary volcanic centers display distinct geochemical characteristics from north to south. Peridotites from the northern section are generally more fer- tile (e.g., Al 2 O 3 = 0.9–4.0%) than those from the south (e.g., Al 2 O 3 = 0.2–2.2%), and have maximum whole-rock Re-depletion Os model ages (T RD ) of 1.8 Ga suggesting their coeval formation in the latest Paleoproterozoic. By contrast, peridotites from the south have maximum T RD model ages that span the Archean–Proterozoic boundary (2.1–2.5 Ga). Peridotites with model ages from both groups are found at Fansi, the southernmost locality in the northern group, which likely marks a litho- spheric boundary. The Neoarchean age of the lithospheric mantle in the southern section matches that of the overlying crust and likely reflects the time of amalgamation of the North China Craton via collision between the Eastern and Western blocks. The Late Paleoproterozoic (1.8 Ga) lithospheric mantle beneath the northern section is significantly younger than the over- lying Archean crust, indicating that the original lithospheric mantle was replaced in this region, either during a major north– south continent–continent collision that occurred during assembly of the Columbia supercontinent at 1.8–1.9 Ga, or from extrusion of 1.9 Ga lithosphere from the Khondalite Belt beneath the northern Trans-North China Orogen, during the 1.85 Ga continental collision between Eastern and Western blocks. Post-Cretaceous heating of the southern section is indi- cated by high temperatures (>1000 °C) recorded in peridotites from the 4 Ma Hebi suite, which are significantly higher than the temperatures recorded in peridotites from the nearby Early Cretaceous Fushan suite (<720 °C), and likely reflects signif- icant lithospheric thinning after the Early Cretaceous. Combining previous Os isotope results on mantle xenoliths from the eastern North China Craton with our new data, it appears that lithospheric thinning and replacement may have evolved from east to west with time, commencing before the Triassic on the eastern edge of the craton, occurring during the Jurassic–Cre- taceous within the interior, and post-dating 125 Ma on the westernmost boundary. Ó 2011 Elsevier Ltd. All rights reserved. 0016-7037/$ - see front matter Ó 2011 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2011.04.018 ⇑ Corresponding author. Tel.: +1 301 825 6189; fax: +1 301 405 3597. E-mail address: [email protected](J. Liu). www.elsevier.com/locate/gca Available online at www.sciencedirect.com Geochimica et Cosmochimica Acta 75 (2011) 3881–3902
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Geochimica et Cosmochimica Acta 75 (2011) 3881–3902
Mapping lithospheric boundaries using Os isotopesof mantle xenoliths: An example from the North China Craton
Jingao Liu a,⇑, Roberta L. Rudnick a, Richard J. Walker a, Shan Gao b, Fu-yuan Wu c,Philip M. Piccoli a, Honglin Yuan d, Wen-liang Xu e, Yi-Gang Xu f
a Department of Geology, University of Maryland, College Park, MD 20742, USAb State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan 430074, China
c State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, P.O. Box 9825,
Beijing 100029, Chinad State Key Laboratory of Continental Dynamics, Department of Geology, Northwest University, Xi’an 710069, China
e College of Earth Sciences, Jilin University, Changchun 130061, Chinaf Key Laboratory of Isotope Geochronology and Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences,
Guangzhou 510640, China
Received 28 December 2010; accepted in revised form 19 April 2011; available 27 April 2011
Abstract
The petrology, mineral compositions, whole rock major/trace element concentrations, including highly siderophile ele-ments, and Re–Os isotopes of 99 peridotite xenoliths from the central North China Craton were determined in order to con-strain the structure and evolution of the deep lithosphere. Samples from seven Early Cretaceous–Tertiary volcanic centersdisplay distinct geochemical characteristics from north to south. Peridotites from the northern section are generally more fer-tile (e.g., Al2O3 = 0.9–4.0%) than those from the south (e.g., Al2O3 = 0.2–2.2%), and have maximum whole-rock Re-depletionOs model ages (TRD) of �1.8 Ga suggesting their coeval formation in the latest Paleoproterozoic. By contrast, peridotitesfrom the south have maximum TRD model ages that span the Archean–Proterozoic boundary (2.1–2.5 Ga). Peridotites withmodel ages from both groups are found at Fansi, the southernmost locality in the northern group, which likely marks a litho-spheric boundary. The Neoarchean age of the lithospheric mantle in the southern section matches that of the overlying crustand likely reflects the time of amalgamation of the North China Craton via collision between the Eastern and Western blocks.The Late Paleoproterozoic (�1.8 Ga) lithospheric mantle beneath the northern section is significantly younger than the over-lying Archean crust, indicating that the original lithospheric mantle was replaced in this region, either during a major north–south continent–continent collision that occurred during assembly of the Columbia supercontinent at �1.8–1.9 Ga, or fromextrusion of �1.9 Ga lithosphere from the Khondalite Belt beneath the northern Trans-North China Orogen, during the�1.85 Ga continental collision between Eastern and Western blocks. Post-Cretaceous heating of the southern section is indi-cated by high temperatures (>1000 �C) recorded in peridotites from the 4 Ma Hebi suite, which are significantly higher thanthe temperatures recorded in peridotites from the nearby Early Cretaceous Fushan suite (<720 �C), and likely reflects signif-icant lithospheric thinning after the Early Cretaceous. Combining previous Os isotope results on mantle xenoliths from theeastern North China Craton with our new data, it appears that lithospheric thinning and replacement may have evolved fromeast to west with time, commencing before the Triassic on the eastern edge of the craton, occurring during the Jurassic–Cre-taceous within the interior, and post-dating 125 Ma on the westernmost boundary.� 2011 Elsevier Ltd. All rights reserved.
0016-7037/$ - see front matter � 2011 Elsevier Ltd. All rights reserved.
3882 J. Liu et al. / Geochimica et Cosmochimica Acta 75 (2011) 3881–3902
1. INTRODUCTION
Xenolithic peridotites transported to the surface by la-vas provide information about the deep lithospheric mantleat the time of eruption. Rhenium–Os isotopic systematicscan potentially date melt depletion events in the peridotitesvia Os model ages (e.g., Walker et al., 1989; Rudnick andWalker, 2009). Assuming that lithosphere formation iscoincident with melt depletion, Os model ages may be usedto determine the age of the lithospheric mantle. In this way,the age of lithospheric mantle can be mapped (e.g., Pearsonet al., 1995) and may provide insights into the tectonicassembly and structure of the continents.
The Late Archean to Paleoproterozoic interval (between2.5 and 1.8 Ga) marks the assembly of the Paleoproterozoic(2.1–1.8 Ga) Columbia supercontinent (e.g., Rogers andSantosh, 2003) and is recorded in the tectonic evolutionof the North China Craton (e.g., Kusky and Li, 2003; Zhaoet al., 2005; Kusky et al., 2007a; Kusky and Santosh, 2009).Understanding this portion of the geologic history of theNorth China Craton provides insights into the configura-tion of the Paleoproterozoic Columbia supercontinent.
However, the history for the Archean–Paleoproterozoicamalgamation of the North China Craton remains contro-versial (e.g., Kusky and Li, 2003; Zhao et al., 2005; Kusky,in press, and references therein). One model suggests thatthe Eastern and Western blocks collided at 1.85 Ga to formthe Trans-North China Orogen, which runs north to southin the central portion of the craton (Fig. 1a), and marks thefinal amalgamation of the North China Craton (e.g., Zhaoet al., 2005). A second model proposes that the Eastern andWestern blocks collided at 2.5 Ga and that this event wasfollowed by a major 1.8–1.9 Ga continent–continent colli-sion along the northern margin during formation of theColumbia supercontinent (Fig. 1b; Kusky and Li, 2003;Kusky et al., 2007a). The critical distinction between thesetwo models lies in the interpretation of a 1.8–1.9 Ga gran-ulite facies metamorphic event in the central North ChinaCraton and a Late Archean ophiolitic complex in the east-ern North China Craton.
Here, we report Re–Os model ages, as well as compre-hensive petrography, major and trace element geochemistryfor 99 xenolithic peridotites (including 67 new data supple-mented by 32 analyses from previous studies; Gao et al.,2002; Becker et al., 2006; Liu et al., 2010) from seven local-ities (i.e., Hannuoba, Yangyuan, Datong, Jining, Fansi,Hebi, and Fushan) covering a broad area in the centralNorth China Craton (Fig. 1). We show that Os isotopesof peridotites have the ability to map deep lithosphericboundaries. Our data provide unique constraints on theLate Archean–Paleoproterozoic tectonic framework of theNorth China Craton, as well as the timing of Mesozoic–Tertiary lithospheric thinning in this region.
2. GEOLOGICAL BACKGROUND
Based on integrated studies of lithology, structure, geo-chronology, and metamorphic pressure–temperature–time(P–T–t) paths, the North China Craton has been dividedinto three main blocks (Fig. 1): the Western Block, the
Eastern Block, and the intervening central region, whichhas been called the Trans-North China Orogen (Fig. 1a;Zhao et al., 2000, 2001, 2005), or the Central Orogenic Belt(Fig. 1b; e.g., Kusky and Li, 2003).
The Western Block is characterized by rather thick(�45 km; Li et al., 2006) Archean crust (Zhao et al.,2001), relatively low surface heat flow (50–60 mW/m2; Huet al., 2000; Tao and Shen, 2008) and thick lithosphere(>150 km; Tian et al., 2009; Chen, 2010). This block hasexperienced only minor Phanerozoic volcanism and rareseismicity. Within the northern portion of the WesternBlock is a nearly east–west trending belt consisting largelyof khondalites (i.e., high-grade metapelitic gneisses com-posed of quartz–feldspar–sillimanite, with graphite, garnetand biotite, ± cordierite) and is, thus, referred to as theKhondalite Belt (Lu et al., 1996) (Fig. 1). P–T–t paths ofthe metamorphic rocks in the belt show isothermal decom-pression suggesting that it formed during a continent–con-tinent collision at ca. 1.9 Ga (Kusky and Li, 2003; Zhaoet al., 2005, 2010; Santosh et al., 2006, 2007; Wan et al.,2006; Dong et al., 2007; Yin et al., 2009).
In contrast to the Western Block, the Eastern Blockmainly consists of relatively thin (30–40 km; Li et al., 2006)Archean crust and thin lithosphere (<100 km; Tian et al.,2009; Chen, 2010), and has relatively high surface heat flow(>64 mW/m2; Hu et al., 2000; Tao and Shen, 2008) and ac-tive seismicity. Studies of xenolithic peridotites in this blockhave suggested that during the Mesozoic, the relatively cold,thick, and refractory lithospheric mantle was removed andreplaced by fertile, thin, Phanerozoic lithospheric mantle,which currently underlies this block (Menzies et al., 1993;Griffin et al., 1998; Gao et al., 2002; Wu et al., 2003, 2006;Rudnick et al., 2004; Zhang et al., 2008; Chu et al., 2009).The westernmost limit of Mesozoic thinning is commonly as-sumed to coincide with the North–South Gravity Lineament(Griffin et al., 1998; Zheng et al., 2001; Menzies et al., 2007;Zhao et al., 2007) that runs through the central portion of theNorth China Craton (Fig. 1).
The roughly north–south trending belt in the centralNorth China Craton (Fig. 1) consists mainly of a series ofNeoarchean to Paleoproterozoic greenschist- to granulite-facies metamorphic terranes (Zhao et al., 2005, and refer-ences therein). The tectonic history of this central regionis in debate. Kusky et al. (2001) suggested that it formedat �2.5 Ga, based on the interpretation of a Late Archeanophiolite complex on the western margin of the EasternBlock (Kusky and Li, 2010), and then experienced granulitefacies metamorphism along the northern boundary in thePaleoproterozoic (ca. 1.85 Ga; Kusky and Li, 2003; Kuskyet al., 2007a), which extends into the Khondalite Belt in theWestern Block (Fig. 1b). Kusky and co-workers refer tothis region as the Central Orogenic Belt. By contrast,“clockwise” P–T paths for the northern granulites thatformed in the Paleoproterozoic have been interpreted tosuggest that the collision between the Eastern and WesternBlocks occurred at ca. 1.85 Ga, forming the so called Trans-North China Orogen (Fig. 1a; e.g., Zhao et al., 2000, 2001;Wilde et al., 2002; Kroner et al., 2005, 2006).
Xenolithic peridotites from Early Cretaceous–Tertiaryvolcanic centers occur over a wide area of the central North
Fig. 1. Tectonic sketch map of the North China Craton composed of the Eastern Block, Western Block, and the central region. In (a), thecentral region is called “Trans-North China Orogen (TNCO)” formed between the Eastern and Western blocks at �1.85 Ga (modified fromZhao et al., 2005). In (b), the central region is called “Central Orogenic Belt (COB)” formed between the Eastern and Western blocks at�2.5 Ga (modified from Kusky and Li, 2003). The Khondalite Belt is a Paleoproterozoic, nearly east–west trending metamorphic belt in theWestern Block that formed earlier than the TNCO (e.g., Zhao et al., 2005; a), while Kusky and Li (2003) suggest that this belt, which formedafter the COB, extends eastwards (b), representing the final amalgamation of the craton. Mantle xenolith localities shown as squares(Paleozoic eruption age), stars (Mesozoic eruption age), and circles (Cenozoic eruption age). The NSGL is the North–South GravityLineament (Griffin et al., 1998). The profile A–A0 is marked for the construction of the age structure of crust and underlying lithosphericmantle in Fig. 9.
Mapping lithospheric boundaries via Os isotopes 3883
China Craton (Fig. 1) and show a large range in composi-tions, from refractory (typical of cratonic peridotites) tofertile (similar to primitive mantle (PM) – a hypotheticalundifferentiated mantle; McDonough and Sun, 1995)(e.g., Zheng et al., 2001; Tang et al., 2008; Xu et al.,2008a, 2010; Liu et al., 2010). The limited age informationavailable for these peridotites (Gao et al., 2002; Zhenget al., 2007; Xu et al., 2008a,b; Zhang et al., 2009; Liuet al., 2010) suggest that they are considerably older thanthe Phanerozoic age of lithospheric mantle sampled byTertiary lavas from the eastern North China Craton(Fig. 1) (Gao et al., 2002; Wu et al., 2003, 2006; Chuet al., 2009), though debate exists about whether the Prote-
rozoic ages reflect mantle formed in the Proterozoic (Gaoet al., 2002; Liu et al., 2010) or refertilized Archean mantle(Tang et al., 2008; Zhang et al., 2009).
3. SAMPLES
The xenoliths studied here come from a wide area in thecentral North China Craton (Fig. 1). Samples from thenorthern section (Hannuoba, Yangyuan, and Datong), aswell as the more southerly Fansi (also spelled “Fanshi”by some authors, e.g., Tang et al., 2008), are all carried inTertiary alkali basalts that erupted to the west of theNorth–South Gravity Lineament (Fig. 1). Samples from
3884 J. Liu et al. / Geochimica et Cosmochimica Acta 75 (2011) 3881–3902
the southern region (Hebi and Fushan), lie to the east ofthis lineament. Fushan hornblende-diorites erupted in theEarly Cretaceous (�125 Ma; Xu et al., 2010), whereas oliv-ine nephelinites at nearby Hebi erupted much later, at 4 Ma(Liu et al., 1990). Jining is located in the Khondalite Belt ofthe Western Block, where xenolith-bearing alkali basaltserupted in the Tertiary period (Zhang and Han, 2006).
The xenoliths investigated here are predominately proto-granular to equigranular, coarse- to medium-grained, gar-net-free spinel lherzolites and harzburgites, as well as raredunites. The modal mineralogy of these samples (ElectronicAnnex EA-1) is illustrated in Fig. 2. Petrography and wholerock compositions of many of these xenolith suites havebeen previously described (Hannuoba: Song and Frey,1989; Chen et al., 2001; Rudnick et al., 2004; Tang et al.,2007; Choi et al., 2008; Zhang et al., 2009; Yangyuan: Xuet al., 2008a; Liu et al., 2010; Fansi: Tang et al., 2008; Hebi:Zheng et al., 2001; and Fushan: Xu et al., 2010). A briefdescription of each suite is summarized in Table 1 and pet-rographic descriptions of the samples investigated here areprovided in Electronic Annex.
In addition to the 32 Re–Os and HSE analyses previ-ously reported for Hannuoba and Yangyuan (Gao et al.,2002; Becker et al., 2006; Liu et al., 2010), we report newelemental and isotopic data for 67 samples (Table 1),including an additional six from Yangyuan, seven from Da-tong, 13 from Jining, 20 from Fansi, 12 from Hebi, and ninefrom Fushan.
4. ANALYTICAL METHODS
4.1. Sample selection and preparation
Mineral mounts of olivine from each of the xenolithscollected at each locality were analyzed in order to assessthe range in degree of melt depletion exhibited by the peri-dotite suites. Olivine compositions were measured using aJEOL 8900 Electron Probe Microanalyzer (EPMA) at theUniversity of Maryland (UMd). The analyses were per-formed using wavelength dispersive spectroscopy with a15 kV accelerating voltage, a 20 nA cup current, and a10 lm diameter beam. A variety of natural minerals were
Olivin
Orthopyroxene
HannuobaYangyuanDatongJiningFansi-Low Fo
HebiFushan
Fansi-High Fo
D
Webste
Olivine we
Harz
burg
ite
Lherzo
Orthopyroxenite
Fig. 2. Petrographic classification of the peridotites, based on proportioHannuoba (Rudnick et al., 2004) and Fushan (Xu et al., 2010).
used as primary and secondary standards. Raw X-rayintensities were corrected using a ZAF algorithm. About1–2 spots per olivine grain, and 2–3 grains of olivines wereanalyzed per sample. Based on the forsterite contents(Fo = molar Mg/(Mg + Fe2+) � 100) of olivines, a repre-sentative suite of peridotites, chosen to span the range inFo contents, was selected from each locality for further ele-mental and isotopic analyses. Portions of the samples werepowdered for whole rock analyses, following the proce-dures detailed by Rudnick et al. (2004).
4.2. Mineral compositions
Major element compositions of olivine, orthopyroxene,clinopyroxene (which may be absent in some harzburgitesand dunites) and spinel were determined on polished thinsections by EPMA at UMd, using the parameters andmethods described above. About two spots per grain,including cores and rims, from three to five grains of eachphase were measured for each sample.
4.3. Whole rock major and trace elements
Whole rock major element compositions were deter-mined by X-ray fluorescence (XRF) on fused glass disksmade from powders (see Boyd and Mertzman (1987) for de-tailed protocols) at Franklin and Marshall College, UnitedStates, or Northwest University, China. Analytical preci-sion and accuracy was typically better than 1% for majorelements of concentrations greater than 0.5% and betterthan 5% for the remaining major elements, as determinedfrom data for international reference rocks analyzed bythese laboratories (e.g., Boyd and Mertzman, 1987;Rudnick et al., 2004). Details regarding the analytical meth-ods for whole rock trace element compositions are providedin Electronic Annex.
4.4. Osmium isotopes and HSE abundances
Mixed 185Re–190Os and HSE (99Ru, 105Pd, 191Ir, 194Pt)spikes were added to each sample powder (1–1.5 g), sealedalong with 3 ml concentrated Teflon distilled HCl and 6 ml
Clinopyroxene
e
10
40
90unite
Clinopyroxeniterite
bsterite
Peridotitelite
Wehrlite
Pyroxenite
ns of olivine and pyroxene. Data sources in addition to this study:
Table 1Petrology of the peridotite suites examined in this study.
Locality n Lithology Size Freshness Sulfidepreservation
Whole rockAl2O3 (%)
Average Fo ofolivine
Hannuoba 16 Lherzolite with rareharzburgite
10–60 cm Fresh Good 1.2–3.8 90.5 ± 0.8
Yangyuan 22 Lherzolite and harzburgite 4–35 cm Fresh Poor 0.9–4.0 90.9 ± 0.6Datong 7 Lherzolite and harzburgite <3 cm Fresh Poor 1.6–3.7 91.0 ± 0.6Jining 13 Lherzolite and harzburgite 4–9 cm but
thinHighly weathered Poor 0.9–6.5 90.4 ± 0.8
Fansi 20 LowFo
Lherzolite with rareharzburgite
3–8 cm Fresh to moderatelyweathered
Poor 1.1–3.9 90.1 ± 0.7
HighFo
Harzburgite 0.8–2.0 92.0 ± 0.3
Hebi 12 Lherzolite and harzburgitewith rare dunite
<4 cm Fresh Good butrare
0.9–2.2 92.0 ± 0.9
Fushan 9 Lherzolite and harzburgitewith rare dunite
3–7 cm Moderatelyweathered
N/A 0.2–1.5 92.0 ± 0.7
Note: n, the number of samples; lithology terms as in Fig. 2; size indicates the maximum length of the specimen; and average Fo of olivine±1r. Data sources provided in text.
Mapping lithospheric boundaries via Os isotopes 3885
concentrated Teflon distilled HNO3 into a chilled, thick-walled borosilicate Carius tube, and heated to 270 �C for4 days. Osmium was extracted from the acid solution usingCCl4 (Cohen and Waters, 1996), then back-extracted intoHBr, and finally purified via microdistillation (Bircket al., 1997). Iridium, Ru, Pt, Pd and Re were separatedand purified using anion exchange column chromatogra-phy, following the steps described in Ireland et al. (2009).
Osmium isotopic measurements were performed by neg-ative thermal ionization mass spectrometry (N-TIMS) atUMd. All samples were measured using a single electronmultiplier on VG Sector 54 or NBS mass spectrometers.Mass fractionation was corrected using 192Os/188Os =3.083, and the internal precision on 187Os/188Os ratioswas typically better than 0.2% (2r). The reported187Os/188Os of samples was corrected for instrumental biasby comparison of the analyzed 187Os/188Os of the Johnson–Matthey Os standard in each analytical session with therecommended value of 0.11380. This correction for187Os/188Os was less than 0.2%.
Using methods identical to this study, Liu et al. (2010)analyzed a subset of whole rock peridotite powders fromXu et al. (2008a) and Zhang et al. (2009), who both utilizedNiS fusion digestion and measured Os using an ICP-MS viasparging (Hassler et al., 2000). A large discrepancy in187Os/188Os (up to 11.6%) and Os concentrations (up to113%) was observed between the two methods (see elec-tronic supplement of Liu et al., 2010, for details). Liuet al. (2010) demonstrated that Carius tube digestion and/or digestion using a high pressure asher, combined withN-TIMS measurement, yielded reproducible Os data forperidotitic reference materials (e.g., UB-N and GP-13) aswell as natural samples. In order to evaluate further thediscrepancies between the Xu/Zhang studies and our data,in particular, whether the generally more radiogenic187Os/188Os values we find result from failure to accessnon-radiogenic Os that may reside within acid-resistantphase, we analyzed two aliquots of relatively refractoryYangyuan sample (YY-22, Fo = 91.2) by the NiS fusion/N-TIMS method. The results are reported in Electronic
Annex (EA-2) and yield 187Os/188Os within uncertainty ofthe previously published results obtained for this samplefrom both high and low temperature Carius tube diges-tion/N-TIMS (Liu et al., 2010). These results demonstratethat the source of the discrepancies is unlikely to reside inthe dissolution method. The reason for the discrepancy re-mains unclear. Since our Os data (and other HSE data)were determined using the same methods validated by Liuet al. (2010), we consider our results robust, and here wedo not include the Re–Os isotope data reported by Xuet al. (2008a) and Zhang et al. (2009) in our discussion.
All other HSE were analyzed using a Nu Plasma MC-ICP-MS at UMd. Isotopic mass fractionation was correctedby periodic measurements of standards (usually one perthree sample analyses) using the standard bracketing meth-od. The accuracy of this analytical method was evaluated bymeasuring reference materials such as UB-N and GP-13 inour laboratory (e.g., Puchtel et al., 2008; Liu et al., 2010),the results of which are comparable, within uncertainties,to those of other labs (e.g., Meisel et al., 2003; Pearsonet al., 2004). Averaged blanks for these measurements areas follows: Os (0.38 ± 0.22 pg), Ir (0.40 ± 0.31 pg), Ru(2.9 ± 2.9 pg), Pt (6.8 ± 2.2 pg), Pd (9.5 ± 3.5 pg) and Re(1.6 ± 0.6 pg) (uncertainties are 1r, calculated from 11blank measurements). Blank corrections for Os, Ir, Ru, Pt,and Pd are negligible (less than 0.2%) for most samples, ex-cept for those with very low HSE concentrations (five Hebisamples (HB-02, HB-09, HB-12, HB-21-2, and HB-22), andthree Fushan samples (FS6-19, FS6-29, and FS6-56), e.g.,having Os generally less than 0.2 ppb), while the Re blankconstitutes 0.3–20% of the total Re in all samples.
5. RESULTS
5.1. Mineral chemistry and equilibration temperatures
The average forsterite contents of olivines analyzed ingrain mounts for xenoliths from all localities (381 xenolithsin total, Electronic Annex EA-3) are plotted as histogramsin Fig. 3. The olivine compositions in the peridotites from
Fig. 3. Histograms of average forsterite contents (Fo of oliv-ine = mol Mg/(Mg + Fe2+)) of olivines from individual peridotitexenoliths and diamond inclusions. (a) Peridotites from Hannuoba(Song and Frey, 1989; Chen et al., 2001; Rudnick et al., 2004; Tanget al., 2007; this study), Yangyuan (Liu et al., 2010; this study), andDatong, which are carried in Tertiary basalts in the northern regionof the central North China Craton, as well as Jining, which lies inthe Khondalite Belt of the Western Block; inset panel (as well as in(b) and (c)) represents the Fo distribution of the samples selectedfor further elemental and isotopic analyses; (b) peridotites fromFansi showing two populations divided into low- and high-Fogroups; (c) peridotites from Fushan (Xu et al., 2010) and Hebi(data for Hebi from this study and Zheng et al., 2001) from thesouthern region; and (d) xenolithic peridotites and olivine inclu-sions in diamonds from Ordovician kimberlites in the easternNorth China Craton (data from Zheng, 1999). The vertical dashedline marks a Fo value of 92.
0 20 40 60 80 100Cr# of spinel
88
90
92
94
Fo o
f oliv
ine
Fansi-H Fo
Fushan
DatongYangyuanHannuoba
Fansi-L Fo
HebiHebi (Zheng)
Jining
Cratonic
OSMA
Abyssal peridotites
FS7-13
more refra
ctory
peridotites
FS2-10
FS-64
FS-50
Fig. 4. Cr# (mol Cr/(Cr + Al) � 100) of spinels vs. Fo of coexistingolivines in peridotite xenoliths from Hannuoba (data from Chenet al., 2001; Rudnick et al., 2004), Yangyuan (data from Liu et al.,2010; this study), Datong, Jining, Fansi (including Fansi L(low)-Foand Fansi H(high)-Fo), Fushan (data from Xu et al., 2010), andHebi (data from Zheng et al., 2001, are shown as gray triangles).Gray field encompasses data for typical cratonic spinel peridotites(Pearson and Wittig, 2008, and references therein). Abyssalperidotites: Arai, 1994, and references therein. OSMA: olivine–spinel mantle array (after Arai, 1994, and references therein).
3886 J. Liu et al. / Geochimica et Cosmochimica Acta 75 (2011) 3881–3902
the northern region (Hannuoba, Yangyuan, and Datong),as well as Jining, are characterized by similar ranges inFo (i.e., 87.4–92.2 with an average of 90.5 ± 0.5 (1r)),reflecting relatively fertile compositions. By comparison,olivines in the peridotites from the southern region (Hebiand Fushan) generally have higher Fo (88.3–92.9 with anaverage of 91.9 ± 0.9; Zheng et al., 2001; Xu et al., 2010;this study) that is similar to, or slightly lower than those
of olivines in cratonic peridotites from the eastern NorthChina Craton (Fig. 3), reflecting the more refractory sub-continental mantle that underlies this region. Fansi perido-tites show a bimodal distribution of Fo contents (Fig. 3) (asin Tang et al., 2008). The low-Fo group (88.0–91.6 with anaverage of 90.1 ± 0.7, 1r) reflects relatively fertile mantle,similar to that present in the other northern localities. Bycontrast, the high-Fo group (91.6–92.5 with an average of92.0 ± 0.3, 1r) represents refractory mantle, similar in com-position to peridotites from Hebi and Fushan in the south.We selected a sub-suite of peridotites from each locality (7–20 samples) that span the observed range of Fo (see insetpanels of Fig. 3), taking into account size and freshness,for further elemental and isotopic analyses.
Major element compositions of minerals of the selectedsamples measured in polished thin sections are given inElectronic Annex EA-4. Olivine compositions analyzed inpolished thin sections are within uncertainties of those mea-sured on mineral mounts. The Cr# (i.e., molar Cr/(Cr + Al) � 100) of spinels from these peridotites correlatewith the Fo contents of olivines; the high-Fo samples gen-erally have higher Cr# (Fig. 4), reflecting melt depletion(Arai, 1994, and references therein). A few samples fall offthe main trend, suggesting disequilibria between olivineand spinel. This includes two samples from Fansi and onefrom Jining from this study, four Hebi samples from thepublished data of Zheng et al. (2001), including the onlythree samples from their “low Mg#” group for which bothspinel and olivine compositions are published, and oneFushan sample from the published data of Xu et al. (2010).
Equilibrium temperatures of the peridotites were calcu-lated using the two-pyroxene thermometer of Brey and
Fig. 5. Whole rock MgO vs. Al2O3 (a) and CaO (b) (in wt%) ofperidotites from Jining (the majority of samples from this suiteshow abnormally high Al2O3), Datong, Hannuoba (Song and Frey,1989; Chen et al., 2001; Rudnick et al., 2004; Choi et al., 2008;Zhang et al., 2009), Yangyuan (Xu et al., 2008a; Liu et al., 2010),Fansi, Fushan (Xu et al., 2010), and Hebi. Symbols as in Fig. 4.Melt depletion typically results in low Al2O3 and CaO and highMgO, which is delineated by the melt depletion trends in the plot.PM: primitive mantle (McDonough and Sun, 1995). Gray fieldshows typical cratonic mantle (Pearson and Wittig, 2008, andreferences therein).
Mapping lithospheric boundaries via Os isotopes 3887
Kohler (1990) and assuming a pressure of 1.5 GPa (close tothe minimum depth (�50 km) of lithospheric mantle, sincepressure cannot be determined in the absence of garnet)(Electronic Annex EA-5). The estimated equilibrium tem-peratures of peridotites from Hannuoba, Yangyuan, Da-tong and Jining are similar at 980–1060, 840–1100, 1060–1100, and 850–990 �C, respectively, although those fromYangyuan and Jining range to lower temperatures than per-idotites from the other localities. The equilibrium tempera-tures of the low-Fo Fansi peridotites (880–1100 �C)overlap with those of the Yangyuan peridotites, while thehigh-Fo Fansi peridotites exhibit a relatively narrow rangeof equilibrium temperatures (960–1040 �C), overlappingthat seen in their low-Fo counterparts. Consistent with pre-vious studies (Zheng et al., 2001), the Hebi peridotites showa narrow range in equilibrium temperatures (1020–1090 �C)that are significantly higher than those of the Fushan peri-dotites (620–720 �C) (Xu et al., 2010). The Fushan perido-tites exhibit the lowest temperatures of all samples. Thereis no correlation between equilibrium temperature and ma-jor element composition in any of the suites.
5.2. Whole rock major and trace elements
Major and/or trace element analyses of xenoliths consid-ered in this study are provided in Electronic Annexes EA-5and EA-6); trace element data are absent for Datong andJining peridotites due to the limited sample powderavailable.
Peridotites from the northern region show overlappingand large ranges in major element compositions and showgood correlations on plots of MgO vs. Al2O3 or CaO, witha few plotting in the field of typical cratonic peridotites(Fig. 5). Such correlations between MgO and Al2O3 orCaO are commonly interpreted to reflect melt depletion(Pearson et al., 2003, and references therein), but can alsobe produced via refertilization (e.g., Le Roux et al., 2007).The majority of the northern peridotites are relatively fertilecompared to cratonic peridotites, consistent with the lowFo of olivine and Cr# of spinel. Jining peridotites have asimilar range of MgO (39.6–45.7%) and CaO (0.4–2.4%)compared to the other northern suites; however theirAl2O3 contents range from 0.9% to higher than the PM(�4.4%; McDonough and Sun, 1995). The abnormally highAl2O3 contents are inconsistent with calculated whole-rockcompositions based on mineral modal abundances (Elec-tronic Annex EA-5), and suggest that such high Al2O3 con-tents are an analytical artifact produced during processingof these rather small and weathered peridotites.
By contrast, peridotites from the southern region aregenerally depleted in Al2O3 (0.9–2.2% and 0.2–1.5% forHebi and Fushan, respectively) and CaO (0.4–2.2% and0.3–1.1%), and are rich in MgO (44–47% and 44–48%)(Xu et al., 2010; this study). Most of them plot in the fieldof cratonic peridotites (Fig. 5).
Xenoliths from Fansi, the southernmost locality in thenorthern region, show a mixed population. The low-Fogroup of Fansi peridotites are compositionally similar(MgO: 37–45%, Al2O3: 1.1–3.9%, and CaO: 0.8–3.4%) tothose from the other northern suites, and distinct from
the high-Fo group of Fansi peridotites (MgO: 44–47%,Al2O3: 0.9–2.0%, and CaO: 0.4–1.4%) that are refractory,like the Hebi and Fushan peridotites.
The data described above reflect a spatial compositionaldichotomy in the lithospheric mantle underlying the centralNorth China Craton. Peridotites from the northern region,including Jining, are relatively fertile, while those from thesouthern region are more refractory in composition, similarto cratonic peridotites. Peridotites from Fansi, located inthe southernmost position within the northern region, in-clude both fertile and refractory compositions.
Whole rock trace element concentrations are provided inElectronic Annex (EA-6 and Figs. EX-1 and EX-2). All per-idotites show depletions of V and Cr, which are a character-istic of partial melting residues. The refractory peridotites
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typically display enriched light rare earth elements (LREE)relative to heavy rare earth elements (HREE), while the rel-atively fertile peridotites show a large variation, rangingfrom depleted to enriched LREE relative to HREE. Moredetailed discussions of these data are provided in ElectronicAnnex.
5.3. Osmium isotopes and HSE abundances
Whole rock Os isotopes and HSE abundances are re-ported in Table 2. For comparison, previously publisheddata from our group for Hannuoba (Gao et al., 2002; Beck-er et al., 2006; Liu et al., 2010) and Yangyuan (Liu et al.,2010), are included.
5.3.1. Hannuoba peridotites
Hannuoba peridotites display high total HSE abun-dances (RHSE = Os + Ir + Ru + Pt + Pd + Re) rangingfrom 17 to 32 ppb, with Ir ranging from 2.5 to 4.3 ppb;Becker et al., 2006; Liu et al., 2010) and patterns that aresimilar to that of model Primitive Upper Mantle (PUM)(Fig. 6a). No obvious Os–Ir fractionation appears in theHannuoba suite. The Pd/Ir ratios ((Pd/Ir)N = 0.7 to 1.5),187Re/188Os (0.045–0.44) and 187Os/188Os (0.116–0.128)are all well correlated with melt depletion indices, such asAl2O3 (Gao et al., 2002; Liu et al., 2010).
5.3.2. Yangyuan peridotites
Yangyuan peridotites have significantly lower HSE con-centrations (RHSE: 3–16 ppb and Ir: 0.7–3.0 ppb) (Liuet al., 2010; this study) compared to Hannuoba peridotites.They are characterized by significant Os, Pd and Re deple-tions, relative to Ir (Fig. 6b; (Pd/Ir)N = 0.11–0.55; Liuet al., 2010), which are not likely to be the result of partialmelting, given their similarity in major element composi-tions to the Hannuoba peridotites. Their 187Os/188Os(0.115–0.126) is fairly well correlated with fertility indices,such as Al2O3 (Liu et al., 2010).
5.3.3. Datong peridotites
Like Yangyuan peridotites, the Datong peridotites arecharacterized by relatively low HSE abundances (RHSE:8–26 ppb and Ir: 1.3–2.3 ppb). Despite the fact that onlyfive samples were analyzed, these peridotites show a diverserange of HSE patterns (Fig. 6c). Sample DAT-15 has aHSE pattern similar to that of PUM. The HSE patternsof DAT-09 and DAT-30 are characterized by Re and plat-inum-like platinum group elements (PPGE: Pt and Pd)depletions relative to iridium-like platinum group elements(IPGE: Os, Ir and Ru). The remaining samples (DAT-05and DAT-31) have variably positive Pt anomalies. Thesesamples, together with the two additional samples analyzedonly for Re–Os isotopes, are characterized by 187Os/188Osranging from 0.115 to 0.126, identical to the range exhibitedby Yangyuan peridotites (Table 2).
5.3.4. Jining peridotites
These are characterized by relatively low and variableHSE abundances (RHSE: 1–16 ppb and Ir: 0.2–3.4 ppb).Large differences are seen in HSE abundances between
some replicates, which are likely due to a nugget effect inthese small samples; however, the HSE pattern shapes forreplicates are similar (Table 2). Two types of HSE patternsare identified (Fig. 6d): one with nearly chondritic Os/Ir ra-tios, and the other with clearly subchondritic Os/Ir ratios.A typical member of the nearly chondritic Os/Ir group ischaracterized by minimal fractionation of the IPGE com-pared to PUM, but exhibits a depletion of Re and PPGE(Pt and Pd). The second group is characterized by con-cave-downward HSE patterns that are similar to those ofYangyuan peridotites (Liu et al., 2010). Jining peridotitesexhibit a large range in 187Os/188Os (0.117–0.128), similarto Hannuoba and Yangyuan peridotite suites.
5.3.5. Fansi peridotites
The HSE abundances of Fansi peridotites are generallylower than those of PUM (RHSE: 8–25 ppb and Ir: 1.5–2.8 ppb). The HSE patterns of these samples are shown inFig. 6e and f for the high-Fo and low-Fo groups, respec-tively. Like Yangyuan, the majority of the Fansi peridotitesare characterized by Re and PPGE depletions relative toIPGE, with a minimal to moderate Os depletion relativeto Ir. Like Datong peridotites DAT-05 and DAT-31, twolow-Fo Fansi peridotites (FS-50 and FS-68) show Pt enrich-ments (Fig. 6f). Sample FS-36 shows a PUM-like HSE pat-tern, similar to DAT-15. In this suite, the high-Fo grouphas low 187Os/188Os (0.110–0.114), whereas the low-Fogroup is characterized by substantially higher 187Os/188Os(0.117–0.127). In spite of a poor correlation between187Re/188Os and 187Os/188Os, a rough positive correlationbetween 187Os/188Os and indicators of fertility (Al2O3 orYb) is present for the majority of the low-Fo samples, butnot for the high-Fo samples, mainly due to a limited rangeof compositions (e.g., Al2O3 = 0.9–1.3 wt%, with one at2.0 wt%).
5.3.6. Hebi peridotites
The Hebi peridotites fall into two groups in terms ofHSE abundances: a low-HSE group (RHSE: 2–7 ppb andIr: 0.1–1.2 ppb), and a high-HSE group (RHSE: 11–67 ppb and Ir: 1.6–7.9 ppb). The patterns of the high-HSE group are generally characterized by chondritic ornear chondritic IPGE, and Re and PPGE depletions rela-tive to the IPGE (Fig. 6g). The low-HSE group is generallycharacterized by positive Ru anomalies (Fig. 6g). Bothgroups have overlapping, low 187Os/188Os (0.112–0.119).Osmium model ages (both rhenium depletion model age,TRD, Walker et al., 1989, and TMA model age, Allegreand Luck, 1980) show similar ranges within the high HSEgroup (TRD = 1.7–2.1 Ga; TMA = 1.8–2.2 Ga; Table 2).The low HSE group shows similar TRD ages (TRD = 1.5–2.0 Ga), but much more variable TMA (TMA = �11 to2.3 Ga) (Table 2), reflecting the more variable Re/Os ratioin the low HSE group.
5.3.7. Fushan peridotites
Like Hebi, the Fushan peridotites have highly variableHSE abundances, and can be divided into two groups:low-HSE (RHSE: 1–2 ppb and Ir: 0.06–0.15 ppb) andhigh-HSE (RHSE: 14–21 ppb and Ir: 2.5–4.5 ppb). The
Note: R represents a replicate analysis of a second aliquot of the same powder. t, eruption time. ‘n.d.’ means ‘not determined’.a Data from Becker et al. (2006) and Liu et al. (2010).b Data from Liu et al. (2010) and this study. Data in italic are from this study.c Fo: forsterite content (molar Mg/(Mg + Fe) � 100).d Cr# of spinel: molar Cr/(Cr + Al) � 100.e The parameters used in model age calculation are kRe = 1.666 � 10–11/year (187Re/188Os)CI = 0.402, (187Os/188Os)CI,0 = 0.1270 (Shirey and Walker, 1998). (187Os/188Os)i represents the initial
value when xenolith was erupted.f PUM: primitive upper mantle. Concentration data from Becker et al. (2006) and Re–Os data from Meisel et al. (2001).
Fig. 6. Primitive-upper-mantle (PUM)-normalized HSE patterns of whole rock peridotites: Hannuoba (a; data from Becker et al., 2006; Liuet al., 2010), Yangyuan (b; data from Liu et al., 2010; this study), Datong (c), Jining (d), Fansi (e, high Fo group; f, low Fo group with somesamples of Pt and/or Pd anomalies), Hebi (g), and Fushan (h). PUM recommended values (Table 2) from Becker et al. (2006).
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high-HSE group is characterized by unfractionated IPGEand strong depletions of Re and PPGE relative to the IPGE(Fig. 6h); it is also characterized by high Os concentrationsof 2.2–4.5 ppb and low 187Os/188Os of 0.110–0.115, whichyield consistent and overlapping model ages (TRD = 1.8–2.5 Ga, TMA = 2.1–2.6 Ga) that are similar to those of theHebi high-HSE peridotites. The patterns of the low-HSEgroup are characterized by positive Ru anomalies(Fig. 6h), like the Hebi low-HSE peridotites, but, unlikeHebi, the low-HSE group (Os = 0.023–0.037 ppb) has moreradiogenic Os isotopic compositions (187Os/188Os = 0.118–0.124). An exceptional sample, dunite FS7-13, has moder-ate HSE abundances (RHSE = 4.4 ppb and Ir = 0.96 ppb)
and is somewhat enriched in Pd relative to Pt. The187Os/188Os of FS7-13 is the only strongly suprachondriticsample (187Os/188Os = 0.261) analyzed from any of thesuites examined here.
6. DISCUSSION
In order to use Os isotopes to establish the age structureof lithospheric mantle beneath the central North ChinaCraton, we need to: (1) determine whether the Re–Os isoto-pic systematics in peridotites have been disturbed by sec-ondary processes, and if so, evaluate their effects on Osmodel ages; and (2) distinguish Proterozoic-aged peridotites
Mapping lithospheric boundaries via Os isotopes 3893
from modern convecting upper mantle, a small fraction ofwhich also yields Proterozoic model ages (e.g., Harveyet al., 2006; Liu et al., 2008; Fig. 7).
6.1. Effects of secondary processes on HSE abundances and
Os isotopic compositions
Secondary processes have clearly affected most of therocks considered here. For example, LREE enrichment isobserved in xenoliths from all suites (e.g., Zheng et al.,2001; Rudnick et al., 2004; Tang et al., 2008; Xu et al.,2008a; Fig. EX-1). The LREE enrichment reflects over-printing and interaction with LREE-enriched melts or flu-ids during one or more events following initial partialmelting of the mantle (Frey and Green, 1974).
Unlike lithophile trace elements, HSE, being both sider-ophile and chalcophile, mainly reside in base metal sulfidesand/or HSE-bearing alloys in mantle rocks (e.g., Hart andRavizza, 1996; Alard et al., 2000; Bockrath et al., 2004;Ballhaus et al., 2006; Lorand et al., 2008, 2010). The degree
Northern region
Freq
uenc
y
0.100 0.110 0.120 0.1300
4
8
12
16
0.140
2.5 0.5 -0.51.53.5
Fansi L Fo
0
2
4
6
8
10a
bAbyssal peridotites Southern region
Fansi H Fo
TRD/Ga
187Os/188Os
187Os/188Os
0
0.25
0.5
0.75
1
0.110 0.120 0.130 0.140
CD
F
Peridotites from northern region (n=63)
Abyssal peridotites worldwide (n=89)
Fig. 7. Histograms of 187Os/188Os of whole rock peridotites fromthe northern region (a, with low Fo Fansi samples plotted), and thesouthern region (b, with high Fo Fansi samples as well as abyssalperidotites plotted). The cumulative distribution functions (CDF)of 187Os/188Os for peridotites from the northern region of theNorth China Craton and abyssal peridotites worldwide are shownas an inset of the upper panel. The CDF for peridotites from thenorthern region plots to the left (i.e., lower values of 187Os/188Os)of that of abyssal peridotites. Abyssal peridotite data are fromParkinson et al. (1998), Brandon et al. (2000), Standish et al.(2002), Harvey et al. (2006), and Liu et al. (2008). The verticaldashed line is at a value of 187Os/188Os = 0.1250.
to which secondary processes have influenced the Re–Osisotopic systematics, and thus, affected the accuracy of theRe–Os chronometer in peridotites can potentially be as-sessed by examining HSE systematics (e.g., Lorand andAlard, 2001; Buchl et al., 2002; Lorand et al., 2004; Reis-berg et al., 2005; Ackerman et al., 2009; Liu et al., 2010).For example, secondary sulfides are typically enriched inPPGE relative to IPGE (e.g., Alard et al., 2000), and theiraddition to residual peridotite should be reflected in en-hanced PPGE and sulfur concentrations in bulk samples(Rudnick and Walker, 2009, and references therein). Below,we discuss the HSE characteristics of xenoliths from eachlocality and explore whether the HSE were significantly im-pacted by secondary processes; we then evaluate possibleimpacts on Os isotopes and model ages.
6.1.1. Hannuoba and Yangyuan
Previous studies have demonstrated that the Os isotopiccompositions of both Hannuoba and Yangyuan peridotiteswere little affected by secondary processes, based mainly onthe following lines of evidence:
(1) In the Hannuoba suite, the observed multiple linearcorrelations between S, Pd/Ir, Re/Os, 187Os/188Os, andimmobile melt depletion indicators such as Al2O3 (Fig. 8),
0.105
0.110
0.115
0.120
0.125
0.130
0 1 2 3 4 5
187 O
s/18
8 Os
Al2O3(Wt. %)
-0.5
3.0
2.0
0.5
1.5
0.0
1.0
2.5
TRD/Ga
0.261
0.7
1.8 Ga
2.5 Ga
PUMFS7-13 FS2-05
FS2-10
FS2-09
DAT-05
DAT-31
FS6-19
Fig. 8. Whole rock Al2O3 vs. 187Os/188Os of peridotites. Themajority of samples from Datong, Hannuoba (Meisel et al., 2001;Gao et al., 2002; Becker et al., 2006; Liu et al., 2010), Yangyuan(Liu et al., 2010; this study), and Fansi low-Fo group follow theupper curve of melt depletions trends, while some outliers aredetailed in the text. Samples from Fushan, Hebi and Fansi high-Fogroup tend to evolve from more ancient mantle sources as outlinedby the lower curve of melt depletion trends. Samples from Jiningare not plotted in this diagram due to abnormal Al2O3 contents.PUM: Meisel et al. (2001) (for 187Os/188Os) and McDonough andSun (1995) (for Al2O3). Symbols as in Fig. 4. The vertical dashedline is Al2O3 of 0.7%, below which it has been proposed thatperidotites lose all Re (Handler et al., 1997). Upper and lowerdashed lines indicate the initial 187Os/188Os of upper and lowercurves of melt depletion trends, respectively. Model TRD ages arecalculated based on a chondritic mantle of 187Os/188Os = 0.127 and187Re/188Os = 0.402 (Shirey and Walker, 1998).
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together with relatively high HSE abundances, primarilyreflect the effects of ancient partial melting, with subsequentgood preservation of primary sulfides and only minor addi-tion of secondary sulfides (Gao et al., 2002; Liu et al., 2010).It is worth emphasizing here that refertilization that signif-icantly post-dates melt depletion would lead to nearhorizontal trends on a plot of Al2O3 vs. 187Os/188Os, asoriginally noted by Reisberg and Lorand (1995) and furtheremphasized by Rudnick and Walker (2009). The fact thatboth the Hannuoba and Yangyuan suites show linear cor-relations having a significant positive slope between PUMand a low 187Os/188Os, low Al2O3 end member on this plot(Fig. EX-3a and b) indicate that these peridotites experi-enced minimal late refertilization (cf. Xu et al., 2008b;Zhang et al., 2009).
(2) HSE characteristics in the Yangyuan suite, such aslow HSE concentrations, Os, Pd and Re depletions relativeto Ir, and low S and Se contents, were interpreted to be aresult of sulfide breakdown in an oxidized environment fol-lowing infiltration of a S-undersaturated melt/fluid (Liuet al., 2010). Because 187Os/188Os correlates positively withfertility indices such as Al2O3 (Fig. 8 and Fig. EX-3b), sul-fide breakdown was interpreted to be a recent phenomenonthat has had little impact on Os isotopic compositions ofthese peridotites (Liu et al., 2010). Thus, the Re–Os isotopicsystematics of Hannuoba and Yangyuan peridotites may beused to constrain their melt depletion ages (Gao et al., 2002;Liu et al., 2010).
6.1.2. Datong
The depletions of Re and PPGE relative to the IPGE insamples DAT-09 and DAT-30 may reflect high degrees ofpartial melting (Pearson et al., 2004), consistent with therefractory compositions of the samples (e.g., Al2O3 = 1.6–2.3%, Fo91.6–91.7). If this interpretation is correct, the non-radiogenic Os isotopic compositions of these samples indi-cate an ancient depletion event. The PUM-like pattern ofsample DAT-15 reflects a low degree of partial melting,consistent with its fertility (e.g., Al2O3 = 3.2%, Fo90.8). Po-sitive Pt anomalies observed in the remaining two samples(DAT-05 and DAT-31) likely reflect the mobility of Pt viamelt percolation involving loss/gain of base metal sulfides(Ackerman et al., 2009).
Collectively, the Datong peridotites show a good posi-tive correlation between 187Os/188Os and 187Re/188Os(r2 = 0.85, with an errorchron age of 1.80 ± 0.56 Ga;Fig. EX-4), and melt depletion indicators, such as Al2O3,although the aforementioned Pt-enriched samples DAT-05 and DAT-31 plot slightly to the right of the 187Os/188Osvs. Al2O3 trend defined by the other samples (Fig. 8). Giventhe HSE fractionation in these two samples, and their rela-tively low Os concentrations, we suggest that their positionto the right of the other Datong samples in Fig. 8 reflectsminor basaltic addition, possibly associated with melt per-colation prior to eruption. Alternatively, the slightly ele-vated Al2O3 contents of these two samples may haveresulted from biased sampling of these very small, rather
coarse-grained peridotites. These observations indicate thatthe growth of 187Os in the Datong peridotites is the result oflong-term decay of radioactive 187Re in a relatively closedresidual system, after varying extents of ancient partialmelting. Evidently melt percolation that may have enrichedAl2O3 did not significantly modify the Re–Os isotopic sys-tematics, consistent with melt–rock mixing trends (Reisbergand Lorand, 1995).
6.1.3. Jining
Those samples with nearly chondritic Os/Ir ratios (sevenout of 13 samples) are characterized by depletions in Reand PPGE relative to the IPGE (Fig. 6d), which is a typicalsignature of a high degree of partial melting (Pearson et al.,2004), consistent with their low 187Os/188Os. By contrast,the concave-downward HSE patterns of samples with lowOs/Ir (Fig. 6d) are similar to those seen in the Yangyuanperidotites, which were previously interpreted to be the re-sult of recent sulfide breakdown after infiltration of an oxi-dative, S-undersaturated melt/fluid (Liu et al., 2010).Although Os was lost relative to Ir, the Os isotopic signa-tures of peridotites were likely little affected by the putativerecent, oxidative sulfide breakdown, as observed in Yangy-uan peridotites (Liu et al., 2010).
6.1.4. Fansi
This suite is characterized by HSE patterns showing Reand PPGE depletions relative to the IPGE, with a minimalto moderate Os depletion relative to Ir (Fig. 6e and f).These patterns are similar to those of the Yangyuan perido-tites interpreted to have experienced recent oxidative sulfidebreakdown through melt percolation (Liu et al., 2010).However, the significant depletions of Re and PPGE rela-tive to the IPGE of the high-Fo samples (Fig. 6e) may haveresulted from high degrees of melting rather than oxidativesulfide breakdown. Their non-radiogenic Os isotopic com-positions must reflect the antiquity of partial melting. ThePUM-like pattern of low-Fo sample FS-36, like DAT-15,reflects a limited degree of melting and little sulfide break-down, consistent with its fertility (e.g., Fo = 89.8 andAl2O3 = 3.5%) and its relatively radiogenic 187Os/188Os(0.1249). Like some Datong peridotites, Pt enrichmentsare present in two low-Fo Fansi samples (FS-50 and FS-68; Fig. 6f), presumably resulting from Pt mobility duringmelt percolation (Ackerman et al., 2009).
The scattered, poor correlations between 187Os/188Osand melt depletion indicators (e.g., Al2O3, Fo, Cr#, and187Re/188Os) for the low-Fo samples of this suite, and forone of the high-Fo samples (F17), presumably point to sig-nificant impacts on either 187Os/188Os, or melt depletionindicators resulting from secondary processes other thanoxidative sulfide breakdown. For example, samples FS-50,FS-64 and FS2-10 show evidence of recent Fe enrichment,either on the basis of olivine that is too Fe-rich relative tocoexisting spinel Cr# (e.g., FS2-10, Fig. 4), or displacementof the samples to the left (low Fo) side of the Fo vs.187Os/188Os trend defined by the other samples (e.g., allthree samples in Fig. EX-5 in Electronic Annex). Moreover,samples FS2-05 and F17 may have experienced addition ofradiogenic Os during melt–rock reaction, given relatively
Mapping lithospheric boundaries via Os isotopes 3895
low Al2O3 (Fig. 8) or high Fo (Fig. EX-5; Table 2) for their187Os/188Os values, respectively; their relatively low Os con-centrations (0.65 and 0.87 ppb, respectively) would makethese samples more susceptible to overprinting. Finally,the slightly elevated Al2O3 contents of FS2-09 and FS2-10may be due to addition of minor amounts of basaltic melt(Liu et al., 2010), which would have had minimal impacton Os isotopic compositions, because basaltic melt nor-mally has 1–3 orders of magnitude lower Os concentrationsthan peridotites (e.g., Walker et al., 1999; Puchtel andHumayun, 2000). In spite of the scatter that is likely gener-ated by secondary processes, the low-Fo Fansi samples dis-play a rough positive correlation between 187Os/188Os andAl2O3 (excluding the outliers discussed above; Fig. 8),which presumably reflects the vestige of ancient partialmelting of these samples.
6.1.5. Hebi
The large range of HSE abundances and patterns in theHebi peridotites may partially reflect the small samplesizes, coupled with inhomogeneous distribution of theHSE-bearing phases, and/or HSE mobility within theupper mantle. Minimal fractionation of IPGE, and Reand PPGE depletions relative to the IPGE (Fig. 6g) inthe high-HSE samples, reflect a high degree of melt extrac-tion (Pearson et al., 2004; Luguet et al., 2007). This is con-sistent with their high Fo values and low Al2O3 contents.By contrast, the low-HSE samples (RHSE: 2–7 ppb) mayhave been influenced by melt percolation (Buchl et al.,2002). However, because melts normally have radiogenicOs isotopic compositions evolving from high Re/Os ratiosand/or radiogenic sources, peridotites stripped of Os dueto melt percolation typically show an enrichment of radio-genic Os isotopic compositions, which is not observed here(187Os/188Os = 0.114–0.119). In theory, total consumptionof HSE-bearing sulfides without formation of HSE-bear-ing alloys at high degrees of melt depletion can lead tosuch low HSE abundances (e.g., Rehkamper et al.,1999), and may account for the low-HSE Hebi peridotites.Such HSE characteristics are also observed in orthopyrox-ene separates from refractory harzburgites that are devoidof visible sulfides and alloys (Luguet et al., 2007). How-ever, it is unclear why some of the Hebi peridotites appar-ently preserved their HSE contents (in refractory sulfidesand/or alloys) at the same degree of melt depletion whileothers did not. Moreover, the positive Ru anomaly ob-served in the low HSE Hebi samples requires at leastone unique stable phase to host Ru, apart from otherHSE in the residues.
The low-HSE Hebi samples generally show chondritic tosuprachondritic Re/Os ratios (187Re/188Os between 0.4 and0.5; Table 2), which, over time, would lead to chondritic tosuprachondritic Os isotopic compositions, instead of theobserved non-radiogenic Os isotopic compositions. Dueto the very low Re concentrations in these samples(0.003–0.010 ppb), a small amount of recent Re addition(e.g., from host basalts) could easily elevate Re/Os ratiosto the observed high values, but would not significantlychange Os isotopic compositions. Overall, the non-radio-genic Os isotopic compositions of both high- and low-
HSE Hebi peridotites must reflect ancient melt depletionexperienced by these rocks.
6.1.6. Fushan
The HSE patterns of the high-HSE Fushan samples arecharacteristic of residues of high degrees of partial melting,whereas the low-HSE samples, like the low-HSE Hebi sam-ples, have experienced depletions of both IPGE and PPGE,with minor Re addition. Further, the two dunites (FS6-19and FS7-13), which have lower Fo contents (90–91) thanthose of lherzolites and harzburgites (�92), were inter-preted to have formed through melt–peridotite reaction(Xu et al., 2010). The suprachondritic 187Os/188Os (0.261)of FS7-13, together with Pd (relative to Pt) enrichment, fur-ther supports dunite formation by melt–peridotite reaction.In the process of melt–peridotite reaction, S-saturated meltsmay precipitate sulfides that control the shape of the HSEpatterns and Os isotopic compositions of the whole rock(Buchl et al., 2002). Compared to FS7-13, sample FS6-19has significantly lower 187Os/188Os (0.1240), but the ratiois still much higher than those of the other two low-HSEsamples FS6-29 and FS6-56 (Table 2). The negative corre-lation between 187Os/188Os and 1/Os (r2 = 0.88) amongthese low-HSE samples (see Fig. EX-6 in Electronic Annex)suggests addition of secondary sulfides, but to a lesser ex-tent for dunite FS6-19, compared to dunite FS7-13. Thus,excluding the dunites that formed by melt–rock reaction,the non-radiogenic Os isotopic compositions of the remain-ing Fushan peridotites document their long-term evolutionunder low Re/Os ratios, reflecting the antiquity of partialmelting.
6.2. Age of the lithospheric mantle
The above discussion suggests that most (but not all) ofthe peridotites studied here have retained their original Osisotope signature with little modification due to secondaryprocesses. The next important consideration before robustmodel ages can be determined is to what degree non-radio-genic Os reflects ancient melting events that led to the for-mation of the continental lithosphere, vs. remnants ofancient melting events that are known to be present as aminor component in the convecting upper mantle (see dis-cussion in Rudnick and Walker, 2009).
When plotted on a histogram (Fig. 7), the 187Os/188Osvalues of peridotites from all suites show a distribution thatis distinct from that of modern abyssal peridotites. Samplesfrom the southern region range to lower 187Os/188Os thanany abyssal peridotite, and their overall distribution, witha peak in 187Os/188Os at �0.114 is clearly resolved frommodern convecting mantle. These samples must, therefore,represent ancient lithospheric mantle. By contrast, therange in 187Os/188Os of peridotites from the northern regioncompletely overlaps that seen in abyssal peridotites. Never-theless, the distribution of 187Os/188Os in these suites is dis-tinct from that seen in abyssal peridotites, or, for thatmatter Mesozoic upper mantle, as sampled by the Jose-phine Ophiolite (Meibom et al., 2002), with the greatestnumber of samples falling below 187Os/188Os = 0.125 (thedotted line). By contrast, though the 187Os/188Os of abyssal
3896 J. Liu et al. / Geochimica et Cosmochimica Acta 75 (2011) 3881–3902
peridotites show the same overall range, the peak in187Os/188Os occurs at more radiogenic values (Fig. 7), con-sistent with their distinct cumulative probability distribu-tions (Fig. 7, inset). Thus, it is highly unlikely that theseperidotites were derived from convecting upper mantle dur-ing the Mesozoic or later. Below we examine the age con-straints provided by the Os isotopic data.
6.2.1. Northern region: Datong, Yangyuan, Hannuoba, and
Jining
Correlations between 187Os/188Os and melt depletionindices (e.g., Al2O3, CaO, or Yb; Handler et al., 1997; Reis-berg et al., 2005) allow the initial 187Os/188Os of a suite ofperidotites to be estimated. Comparison of this ratio witha model for Os isotopic evolution in the mantle, allows der-ivation of a model age for the suite of peridotites. The187Os/188Os values of peridotites from the northern regionfor which we have reliable Al2O3 concentrations (i.e., Han-nuoba, Datong and Yangyuan) generally show positive cor-relations with Al2O3 (Fig. 8). Using linear regression of thedata and extrapolating to 0.7% Al2O3 (suggested to be theAl2O3 value at or below which all Re is removed in perido-tites during progressive melt depletion (Handler et al.,1997)), we estimate the initial 187Os/188Os of the three north-ern peridotite suites to be�0.115. This isotopic compositioncorresponds to an Os model age of ca. 1.8 Ga (Fig. 8), whichis similar to the maximum TRD ages of peridotites from eachlocality (Table 2). This is generally consistent with the oldestPaleoproterozoic TRD (�1.5 Ga) and TMA (1.8–2.0 Ga) agesof sulfides (187Re/188Os < 0.1) from Hannuoba peridotites(Xu et al., 2008b), and a Lu–Hf errorchron age (1.7 ±0.1 Ga; 2r) of clinopyroxenes separated from Yangyuanperidotites (Liu et al., in press). Although the Jining perido-tites show no good correlation between 187Os/188Os andAl2O3, mainly due to their abnormal Al2O3 contents, their187Os/188Os ratios do correlate negatively with olivine Focontent and the Cr# of spinels (see Fig. EX-5 in ElectronicAnnex), as would be expected if they formed as residues ofearlier partial melting. As their 187Os/188Os ratios (0.117–0.128) overlap with peridotites from the northern region(0.115–0.128), the Jining peridotites probably experiencedmelt extraction at the same time, e.g., about 1.8 Ga ago.We conclude that the lithosphere beneath the northern re-gion formed by melt depletion at ca. 1.8 Ga. This is signifi-cantly younger than the age of crust formation in thisregion, which is 2.4–2.8 Ga (Wu et al., 2005a).
6.2.2. Southern region: Hebi and Fushan
In comparison with peridotites in the northern region,peridotites from the southern region (Hebi and Fushan)do not exhibit good correlations between 187Os/188Os andAl2O3, largely due to the lack of a spread in Al2O3 resultingfrom their generally refractory compositions. For refrac-tory peridotites, with low Re/Os ratios, the TRD modelage should approximate the timing of melt depletion. Inaddition, if Re/Os has not been affected by secondary pro-cesses, the TMA Os model age may more accurately repre-sent the melt depletion age than TRD age. Low Re/Osratio samples will show little difference between TRD andTMA.
As shown above, the low-HSE samples in both Hebi andFushan are more susceptible to overprinting. We thereforefocus our attention here on the high-HSE samples in orderto constrain the timing of melt depletion and, hence, litho-spheric formation. These high-HSE peridotites are charac-terized by low 187Os/188Os of 0.112–0.115 (correspondingto TRD = 1.7–2.1 Ga and TMA = 1.8–2.3 Ga) for Hebiand 0.110–0.115 (TRD = 1.8–2.5 Ga and TMA = 2.1–2.6 Ga) for Fushan, respectively. The small differences(<0.3 Ga) between TRD and TMA ages are consistent withtheir low Re/Os ratios, which are characteristic of perido-tites that experienced high degrees of partial melting. Theancient TRD and TMA ages suggest that peridotites fromboth Hebi and Fushan experienced melt depletion event(s)in the Neoarchean to Paleoproterozoic (2.1–2.5 Ga). Theseages are equal to or slightly younger than the ages derivedfrom in situ Re–Os analyses reported for two sulfide grainsfrom two Hebi peridotites (Zheng et al., 2007); the range toolder ages in the sulfides (with TMA of 2.5 Ga and 3.0 Ga,respectively) may reflect isotopic heterogeneity of the man-tle that melted at �2.5 Ga, or minor overprinting of thewhole rocks due to sulfide metasomatism, though no sec-ondary sulfides have been observed here or described inprevious studies (Zheng et al., 2007). We conclude that per-idotites from the southern region formed by melt depletionat �2.5 Ga, which is earlier than those from the northernregion (ca. 1.8 Ga), but similar to the age of the crust in thisregion (Wu et al., 2005a; Liu et al., 2009).
6.2.3. Fansi
The high-Fo samples have non-radiogenic 187Os/188Osof 0.110–0.114, corresponding to Neoarchean to Paleopro-terozoic model ages (TRD = 2.0–2.5 Ga and TMA = 2.2–2.7 Ga), similar to those of the Hebi and Fushan perido-tites. By contrast, the low-Fo samples display a range ofhigher 187Os/188Os (0.117–0.126) that largely overlaps withthe range seen in other peridotites from the northern region(0.115–0.128) (Figs. 7 and 8). Excluding the few samplesthat show overprinting of either 187Os/188Os or Al2O3 fromsecondary processes, there appears to be a crude positivecorrelation between 187Os/188Os and Al2O3. Using a linearregression of these data, we derive a formation model ageof �1.8 Ga at 0.7% Al2O3, identical to that of peridotitesfrom the northern region (Fig. 8). The low-Fo and high-Fo samples have overlapping Al2O3 contents (1.1–3.9%vs. 0.9–2.0%, respectively), but there is a large differencein 187Os/188Os values (0.117–0.126 vs. 0.110–0.114) at a gi-ven Al2O3. This observation demonstrates that the low-Fosamples were not derived from the high-Fo samples by meltaddition or melt–peridotite reaction (cf. Tang et al., 2008),which would have led to Al2O3 enrichment, while havinglittle impact on 187Os/188Os (e.g., Reisberg and Lorand,1995; Rudnick and Walker, 2009, and references therein;Liu et al., 2010). Therefore, we conclude that there aretwo ages of lithospheric mantle beneath Fansi: refractorymantle that underwent melt depletion in the Neoarchean–Paleoproterozoic (�2.2–2.5 Ga), like peridotites from thesouthern region, and more fertile mantle that underwentmelt depletion in the late Paleoproterozoic (�1.8 Ga). Noclear correlation is observed between calculated equilibrium
Mapping lithospheric boundaries via Os isotopes 3897
temperatures and ages, suggesting that ancient, refractory,and younger, fertile mantle may not be stacked upon oneanother, but are probably interleaved.
6.3. Tectonic implications
The age constraints discussed above document two agesof lithospheric mantle beneath the central North ChinaCraton: at �1.8 Ga in the north, and at �2.5 Ga in thesouth, with a boundary between these two age provincesrunning through the Fansi locality (Figs. 1 and 9). Theseages provide information about the tectonic history of thissection of the North China Craton and, combined withequilibrium temperatures, the timing of lithospheric thin-ning beneath this portion of the craton.
6.3.1. Precambrian tectonics in central North China Craton
The similarity between the age of the crust and litho-spheric mantle in the southern region (Neoarchean to EarlyPaleoproterozoic) suggests that �2.5 Ga represents the timeof cratonization of this portion of the North China Craton(Wu et al., 2005a; Liu et al., 2009). By contrast, the decid-edly younger lithospheric mantle underlying Archean crustin the northern region points to replacement of the originalArchean lithospheric mantle at �1.8 Ga. Several mecha-nisms have been advanced to explain younger lithosphericmantle underlying older continental crust, including ther-mal/chemical erosion (e.g., Griffin et al., 1998; Zhenget al., 2001), density foundering (often referred to as“delamination”, (Kay and Kay, 1993; Lee et al., 2000;Gao et al., 2002; Wu et al., 2006), lateral escape during col-lision (Menzies et al., 1993; Wu et al., 2006), and transfor-mation through refertilization (e.g., Zhang et al., 2002,2009; Zhang, 2005). Of these, only density foundering canexplain removal of the entire Archean lithospheric mantleand its replacement by juvenile mantle in the Paleoprotero-zoic; the other mechanisms would presumably predict pres-ervation of relict Archean lithospheric mantle at shallowdepths beneath the region, which is not observed in thexenolith suites.
Previous studies of metamorphic crustal rocks have doc-umented widespread regional granulite facies metamor-phism at �1.8–1.9 Ga in the northern region as well asthe Khondalite Belt (e.g., Zhao et al., 2005, and referencestherein). These metamorphic rocks record clockwise P–T
paths that are interpreted to have resulted from conti-nent–continent collisional event(s). Thus, the �1.8 Ga
Hebi Fushan Fansi DT YY JN HNB
Crust
LithosphericMantle
N
~2.5 Ga
~2.5 Ga ~1.8 Gaboth groups
A A’
Fig. 9. The age dichotomy of crust and lithospheric mantlebeneath the central North China Craton along the profile A–A0
(see Fig. 1). HNB: Hannuoba; JN: Jining; YY: Yangyuan; and DT:Datong.
lithospheric mantle replacement could be the result of the�1.8–1.9 Ga continent–continent collision(s). However, itis currently debated how the collisional processes proceededin the tectonic framework of the North China Craton (e.g.,Kusky and Li, 2003; Zhai and Liu, 2003; Zhao et al., 2005;Kusky, in press).
Zhao and co-workers suggest that there were two sepa-rate collisional events in this region: one at �1.95 Ga result-ing in the formation of the Khondalite Belt, which marksthe amalgamation of the Western Block, and a second at�1.85 Ga, when the Western and Eastern blocks collidedto form the Trans-North China Orogen, representing the fi-nal assembly of the North China Craton (Zhao et al., 2005,and references therein; Fig. 1a). However, this tectonic sce-nario does not explain the north–south age dichotomy ofthe lithospheric mantle documented here (Fig. 9). Forexample, why would the lithospheric mantle replacementoccur only in the northern region, not throughout the entireTrans-North China Orogen? One possible explanation isthat lithospheric replacement was caused by greater short-ening in the north than in the south of the Trans-NorthChina Orogen. This appears to be broadly consistent withthe current geometry of the Trans-North China Orogen,with the northern section narrower than the southern(Fig. 1a). Another possibility is that the Late Paleoprotero-zoic age of lithospheric mantle in the north was the result ofextrusion of �1.9 Ga lithosphere from the Khondalite Beltbeneath the northern Trans-North China Orogen duringthe �1.85 Ga collision of the Western and Eastern blocks(Fig. 1a), assuming that the Khondalite belt formed earlierthan the Trans-North China Orogen (e.g., Yin et al., 2009;Zhao et al., 2010).
Alternatively, Kusky et al. (2001, 2007b) suggest that thecentral North China Craton marks the site of collision ofthe Eastern and Western blocks at �2.5 Ga. The craton la-ter experienced a major �1.8–1.9 Ga continent–continentcollision event along the northern margin during amalgam-ation of the Columbia supercontinent (Kusky, in press, andreferences therein). In the Kusky model, the KhondaliteBelt constitutes part of the Paleoproterozoic continent–con-tinent collision in the northern North China Craton(Fig. 1b). In this scenario, the observed north–south agedichotomy of lithospheric mantle would reflect removal ofthe original Neoarchean lithospheric mantle and replace-ment of juvenile mantle along the northern margin of thecraton due to the 1.8–1.9 Ga continent–continent collision,while the lithospheric mantle in the southern region reflectsits preservation since it formed during the �2.5 Ga collisionbetween the Eastern and Western blocks of the craton.
In summary, any model that seeks to explain the assem-bly history of the North China Craton should be able topredict the age dichotomy of lithospheric mantle docu-mented here. While these data cannot currently be usedto eliminate either tectonic model for the central NorthChina Craton, aspects of the Kusky model would seem tofit the observations with little special pleading. In this sce-nario, Fansi would mark the northern boundary of litho-spheric replacement that occurred due to collision alongthe northern margin of the North China Craton duringthe Paleoproterozoic assembly of the Columbia superconti-
3898 J. Liu et al. / Geochimica et Cosmochimica Acta 75 (2011) 3881–3902
nent. Alternatively, if the Zhao model is correct, the agedichotomy of lithospheric mantle in the central regionwould require either significant differences in the degree oflithospheric shortening from north to south in the Trans-North China Orogen, resulting in lithospheric removal inthe north, but not the south, or extrusion of �1.9 Ga lith-osphere from the Khondalite Belt beneath the northernTrans-North China Orogen during the �1.85 Ga continen-tal collision between Eastern and Western blocks, with Fan-si marking the southern boundary.
6.3.2. Timing of Phanerozoic lithospheric thinning
Lithospheric thinning beneath the eastern North ChinaCraton occurred largely in the Mesozoic, as reflected bythe timing of the magmatic flare-up that occurred through-out the Eastern Block (e.g., Xu, 2001; Wu et al., 2005b) andthe change in lithospheric mantle composition reflected byxenolithic peridotites carried by Ordovician kimberlitesand those in Tertiary alkali basalts (e.g., Menzies et al.,1993; Griffin et al., 1998; Chu et al., 2009). However, the ex-act timing, geometry and vertical and lateral extent of thin-ning/replacement is still emerging, largely from studies ofmantle xenoliths.
Mantle xenoliths found in a few Mesozoic localities pro-vide important insights into the spatial and temporal signa-ture of the thinning. For example, peridotites from theeastern edge of the North China Craton hosted by TriassicKorean kimberlites (Fig. 1) are characterized by elementaland Os isotopic compositions similar to modern convectivemantle (Yang et al., 2010), suggesting that the lithosphericremoval and replacement occurred no later than the Trias-sic beneath this region. Xenoliths in the Early CretaceousFuxin (�100 Ma) and Laiwu (125–133 Ma) localities, bothof which occur in the center of the Eastern Block (Fig. 1),sample fragments of refractory, Archean lithospheric man-tle but also appear to contain a proportion of fertile perido-tites (Zheng et al., 2007; Gao et al., 2008), which can beinterpreted to suggest that the mantle replacement wason-going during that time. In addition, unusual composi-tions of Mesozoic magmas in this region suggest that den-sity foundering was occurring beneath the Eastern Blockduring the Late Jurassic to Early Cretaceous (Gao et al.,2004, 2008).
The data presented here shed additional light on the tim-ing of lithospheric thinning beneath the North China Cra-ton. Xenolithic samples from Fushan, which erupted at125 Ma on the western boundary of the Eastern Block,are predominantly composed of highly refractory perido-tites with Neoarchean to Early Paleoproterozoic Os modelages. This, together with very low estimated equilibriumtemperatures (620–720 �C), suggests the presence of thick,cold, refractory Archean lithospheric mantle, which, inturn, indicates that lithospheric thinning probably com-menced after 125 Ma in this region.
By contrast, peridotites from the nearby 4 Ma Hebilocality are also fragments of refractory Archean litho-spheric mantle (Zheng et al., 2001, 2007; this study), butthey have significantly higher equilibrium temperatures(1020–1090 �C). The lack of garnet in these peridotites lim-its their equilibrium depths to <100 km (O’Neill, 1981).
Thus, the contrast in equilibrium temperatures betweenFushan and Hebi presumably reflects an increase in thegeotherm associated with lithospheric thinning, implyingthat complete lithospheric removal did not occur in this re-gion, but the lithosphere was significantly thinned after theEarly Cretaceous. Combining all observations, a picture be-gins to emerge of lithospheric thinning/removal commenc-ing from east (Triassic, Yang et al., 2010) to west(<125 Ma, this study) in the North China Craton duringthe Mesozoic, with complete removal of the ancientlithospheric mantle in the east, and partial removal in thewest.
7. CONCLUSIONS
The data reported here for nearly 100 peridotitic xeno-liths from seven suites within the central North China Cra-ton allow us to map lithospheric boundaries in this region.These data, in turn, shed light on the Precambrian accretionhistory of the craton, as well as Mesozoic to Tertiarythinning.
(1) Peridotites from the northern region are generallymore fertile than those from the south, and havemaximum TRD model ages suggesting their coevalformation at �1.8 Ga. By contrast, peridotites fromthe southern region have older (2.1–2.5 Ga) maxi-mum TRD model ages. Peridotites with model agesof both groups are found at Fansi. Thus, therewas diachronous formation of lithospheric mantlefrom north to south, with the boundary at or nearFansi.
(2) Crust and lithospheric mantle have the same age inthe southern region, whereas the lithospheric mantleis significantly younger than the overlying crust inthe northern region. The coupled Neoarchean crustand mantle in the southern region marks the timeof cratonization in this region. The crust–mantledecoupling in the north documents lithosphericmantle replacement at �1.8 Ga, likely resultingeither from a �1.8–1.9 Ga continent–continent colli-sion associated with amalgamation of the Columbiasupercontinent, or from a large difference in thedegree of tectonic shortening from north (more) tosouth (less), or extrusion of �1.9 Ga lithospherefrom the Khondalite Belt beneath the northernTrans-North China Orogen, during the �1.85 Gacontinental collision between Eastern and Westernblocks.
(3) The age structure of lithospheric mantle beneath theNorth China Craton recorded in mantle xenolithserupted from the Paleozoic, through Mesozoic toCenozoic, suggests that lithospheric thinning andreplacement may have evolved from east to west withtime, starting with mantle lithosphere removal beforethe Triassic on the eastern edge of the craton in theKorean peninsula, occurring during the Jurassic–Cretaceous within the interior of the Eastern Block,to thinning that post-dates 125 Ma on the western-most boundary of the Eastern Block.
Mapping lithospheric boundaries via Os isotopes 3899
ACKNOWLEDGMENTS
Igor Puchtel, Richard Ash, and Lynnette Pitcher are thankedfor help in chemical separation and/or mass spectrometric mea-surements. Evan Smith is acknowledged for analyzing olivine forsome of the Hannuoba peridotites. Wei Yang, Jeremy Bellucci,Sara Peek, James Day and Tom Ireland are thanked for readingand improving earlier versions of this manuscript. We also appre-ciate the support from the Maryland NanoCenter and the NispLabthat houses the EPMA. The NispLab is supported in part by theNSF as a MRSEC shared experimental facility. This work was sup-ported by U.S. NSF (Grants EAR 0635671 and 0911096 to R.L.R.and R.J.W.), and the National Nature Science Foundation of Chi-na (Grants 40821061, 90714010, 90814003), as well as the Ministryof Education of China (B07039). We thank John Lassiter, TimothyKusky, Guochuan Zhao and an anonymous reviewer for their com-prehensive and constructive comments that helped to strengthenthe paper. Martin Menzies is thanked for efficient editorialhandling.
APPENDIX A. SUPPLEMENTARY DATA
Supplementary data associated with this article can befound, in the online version, at doi:10.1016/j.gca.2011.04.018.
REFERENCES
Ackerman L., Walker R. J., Puchtel I. S., Pitcher L., Jelinek E. andStrnad L. (2009) Effects of melt percolation on highly sidero-phile elements and Os isotopes in subcontinental lithosphericmantle: a study of the upper mantle profile beneath CentralEurope. Geochim. Cosmochim. Acta 73, 2400–2414.
Alard O., Griffin W. L., Lorand J. P., Jackson S. E. and O’Reilly S.Y. (2000) Non-chondritic distribution of the highly siderophileelements in mantle sulphides. Nature 407, 891–894.
Allegre C. J. and Luck J. M. (1980) Osmium isotopes aspetrogenetic and geological tracers. Earth Planet. Sci. Lett.
48, 148–154.
Arai S. (1994) Characterization of spinel peridotites by olivinespinel compositional relationships – review and interpretation.Chem. Geol. 113, 191–204.
Ballhaus C., Bockrath C., Wohlgemuth-Ueberwasser C., LaurenzV. and Berndt J. (2006) Fractionation of the noble metals byphysical processes. Contrib. Mineral. Petrol. 152, 667–684.
Becker H., Horan M. F., Walker R. J., Gao S., Lorand J. P. andRudnick R. L. (2006) Highly siderophile element compositionof the Earth’s primitive upper mantle: constraints from newdata on peridotite massifs and xenoliths. Geochim. Cosmochim.
Acta 70, 4528–4550.
Birck J. L., RoyBarman M. and Capmas F. (1997) Re–Os isotopicmeasurements at the femtomole level in natural samples.Geostand. Newsl: J. Geostand. Geoanal. 21, 19–27.
Bockrath C., Ballhaus C. and Holzheid A. (2004) Fractionation ofthe platinum-group elements during mantle melting. Science
305, 1951–1953.
Boyd F. R. and Mertzman S. A. (1987) Composition of structure ofthe Kaapvaal lithosphere, southern Africa. In Magmatic
Processes – Physicochemical Principles, vol. 1 (ed. B. O.Mysen). The Geochemical Society, Special Publication.
Brandon A. D., Snow J. E., Walker R. J., Morgan J. W. and MockT. D. (2000) Pt-190–Os-186 and Re-187–Os-187 systematics ofabyssal peridotites. Earth Planet. Sci. Lett. 177, 319–335.
Brey G. P. and Kohler T. (1990) Geothermobarometry in four-phase lherzolites: II. New thermobarometers, and practicalassessment of existing thermobarometers. J. Petrol. 31, 1353–
1378.
Buchl A., Brugmann G., Batanova V. G., Munker C. andHofmann A. W. (2002) Melt percolation monitored by Osisotopes and HSE abundances: a case study from the mantlesection of the Troodos Ophiolite. Earth Planet. Sci. Lett. 204,
385–402.
Chen L. (2010) Concordant structural variations from the surfaceto the base of the upper mantle in the North China Craton andits tectonic implications. Lithos 120, 96–115.
Chen S. H., O’Reilly S. Y., Zhou X. H., Griffin W. L., Zhang G.H., Sun M., Feng J. L. and Zhang M. (2001) Thermal andpetrological structure of the lithosphere beneath Hannuoba,Sino-Korean Craton, China: evidence from xenoliths. Lithos
56, 267–301.
Choi S. H., Mukasa S. B., Zhou X. H., Xian X. H. and AndronikoA. V. (2008) Mantle dynamics beneath East Asia constrainedby Sr, Nd, Pb and Hf isotopic systematics of ultramaficxenoliths and their host basalts from Hannuoba, North China.Chem. Geol. 248, 40–61.
Chu Z. Y., Wu F. Y., Walker R. J., Rudnick R. L., Pitcher L.,Puchtel I. S., Yang Y. H. and Wilde S. A. (2009) Temporalevolution of the lithospheric mantle beneath the eastern NorthChina Craton. J. Petrol. 50, 1857–1898.
Cohen A. S. and Waters F. G. (1996) Separation of osmium fromgeological materials by solvent extraction for analysis bythermal ionisation mass spectrometry. Anal. Chim. Acta 332,
269–275.
Dong C. Y., Liu D. Y., Li J. J., Wan Y. S., Zhou H. Y., Li C. D.,Yang Y. H. and Xie L. W. (2007) Palaeoproterozoic Khonda-lite Belt in the western North China Craton: new evidence fromSHRIMP dating and Hf isotope composition of zircons frommetamorphic rocks in the Bayan Ul-Helan Mountains area.Chin. Sci. Bull. 52, 2984–2994.
Frey F. A. and Green D. H. (1974) Mineralogy, geochemistry andorigin of lherzolite inclusions in Victorian basanites. Geochim.
Cosmochim. Acta 38, 1023–1059.
Gao S., Rudnick R. L., Carlson R. W., McDonough W. F. and LiuY. S. (2002) Re–Os evidence for replacement of ancient mantlelithosphere beneath the North China craton. Earth Planet. Sci.
Lett. 198, 307–322.
Gao S., Rudnick R. L., Xu W. L., Yuan H. L., Liu Y. S., WalkerR. J., Puchtel I. S., Liu X. M., Huang H., Wang X. R. andYang J. (2008) Recycling deep cratonic lithosphere andgeneration of intraplate magmatism in the North China Craton.Earth Planet. Sci. Lett. 270, 41–53.
Gao S., Rudnick R. L., Yuan H. L., Liu X. M., Liu Y. S., Xu W.L., Ling W. L., Ayers J., Wang X. C. and Wang Q. H. (2004)Recycling lower continental crust in the North China craton.Nature 432, 892–897.
Griffin W. L., Zhang A. D., O’Reilly S. Y. and Ryan C. G. (1998)Phanerozoic evolution of the lithosphere beneath the Sino-Korean Craton. In Mantle Dynamics and Plate Interactions in
East Asia, vol. 27 (eds. M. Flower, S.-L. Chung, C.-H. Lo andT.-Y. Lee). American Geophysical Union, Washington, DC.Geodynamics Series, pp. 107–126.
Handler M. R., Bennett V. C. and Esat T. M. (1997) Thepersistence of off-cratonic lithospheric mantle: Os isotopicsystematics of variably metasomatised southeast Australianxenoliths. Earth Planet. Sci. Lett. 151, 61–75.
Hart S. R. and Ravizza G. (1996) Os partitioning between phasesin lherzolite and basalt. Isotopic studies of crust–mantleevolution. AGU Monogr. 95, 123–134.
3900 J. Liu et al. / Geochimica et Cosmochimica Acta 75 (2011) 3881–3902
Harvey J., Gannoun A., Burton K. W., Rogers N. W., Alard O.and Parkinson I. J. (2006) Ancient melt extraction from theoceanic upper mantle revealed by Re–Os isotopes in abyssalperidotites from the Mid-Atlantic ridge. Earth Planet. Sci. Lett.
244, 606–621.
Hassler D. R., Peucker-Ehrenbrink B. and Ravizza G. E. (2000)Rapid determination of Os isotopic composition by spargingOsO4 into a magnetic-sector ICP-MS. Chem. Geol. 166, 1–14.
Hu S. B., He L. J. and Wang J. Y. (2000) Heat flow in thecontinental area of China: a new data set. Earth Planet. Sci.
Lett. 179, 407–419.
Ireland T. J., Walker R. J. and Garcia M. O. (2009) Highlysiderophile element and Os-187 isotope systematics of Hawai-ian picrites: implications for parental melt composition andsource heterogeneity. Chem. Geol. 260, 112–128.
Kay R. W. and Kay S. M. (1993) Delamination and delaminationmagmatism. Tectonophysics 219, 177–189.
Kroner A., Wilde S. A., Li J. H. and Wang K. Y. (2005) Age andevolution of a late Archean to Paleoproterozoic upper to lowercrustal section in the Wutaishan/Hengshan/Fuping terrain ofnorthern China. J. Asian Earth Sci. 24, 577–595.
Kroner A., Wilde S. A., Zhao G. C., O’Brien P. J., Sun M., Liu D.Y., Wan Y. S., Liu S. W. and Guo J. H. (2006) Zircongeochronology and metamorphic evolution of mafic dykes inthe Hengshan Complex of northern China: evidence for latePalaeoproterozoic extension and subsequent high-pressuremetamorphism in the North China Craton. Precambrian Res.
146, 45–67.
Kusky T. M., Li J. H. and Tucker R. D. (2001) The ArcheanDongwanzi ophiolite complex, North China craton: 2.505-billion-year-old oceanic crust and mantle. Science 292, 1142–
1145.
Kusky T. M. and Li J. H. (2003) Paleoproterozoic tectonicevolution of the North China Craton. J. Asian Earth Sci. 22,
383–397.
Kusky T., Li J. G. and Santosh M. (2007a) The PaleoproterozoicNorth Hebei Orogen: North China craton’s collisional suturewith the Columbia supercontinent. Gondwana Res. 12, 4–28.
Kusky T. M., Zhi X. C., Li J. H., Xia Q. X., Raharimahefa T. andHuang X. N. (2007b) Chondritic osmium isotopic compositionof Archean ophiolitic mantle, North China craton. Gondwana
Res. 12, 67–76.
Kusky T. M. and Santosh M. (2009) The Columbia connection inNorth China. In Paleoproterozoic Supercontinents and Global
Evolution, vol. 323 (eds. S. M. Reddy, R. Mazumder, D. Evansand A. S. Collins). Geological Society, Special Publication,
London, pp. 49–71.
Kusky T. M. and Li J. H. (2010) Origin and emplacement ofArchean ophiolites of the Central Orogenic Belt, North ChinaCraton. J. Earth Sci. 21, 744–781.
Kusky T. M. (in press) Geophysical and geological tests of tectonicmodels of the North China Craton. Gondwana Res.
Le Roux V., Bodinier J. L., Tommasi A., Alard O., Dautria J. M.,Vauchez A. and Riches A. J. V. (2007) The Lherz spinellherzolite: refertilized rather than pristine mantle. Earth Planet.
Sci. Lett. 259, 599–612.
Lee C. T., Yin Q. Z., Rudnick R. L., Chesley J. T. and Jacobsen S.B. (2000) Osmium isotopic evidence for mesozoic removal oflithospheric mantle beneath the Sierra Nevada, California.Science 289, 1912–1916.
Li C., van der Hilst R. D. and Toksoz A. N. (2006) Constraining P-wave velocity variations in the upper mantle beneath SoutheastAsia. Phys. Earth Planet. In. 154, 180–195.
Liu C. Z., Snow J. E., Hellebrand E., Brugmann G., von der HandtA., Buchl A. and Hofmann A. W. (2008) Ancient, highly
Liu D. Y., Wilde S. A., Wan Y. S., Wang S. Y., Valley J. W., KitaN., Dong C. Y., Xie H. Q., Yang C. X., Zhang Y. X. and GaoL. Z. (2009) Combined U–Pb, hafnium and oxygen isotopeanalysis of zircons from meta-igneous rocks in the southernNorth China Craton reveal multiple events in the LateMesoarchean–Early Neoarchean. Chem. Geol. 261, 139–153.
Liu J. G., Rudnick R. L., Walker R. J., Gao S., Wu F. Y. andPiccoli P. M. (2010) Processes controlling highly siderophileelement fractionations in xenolithic peridotites and theirinfluence on Os isotopes. Earth Planet. Sci. Lett. 297, 287–297.
Liu J. G., Carlson R. W., Rudnick R. L., Walker R. J., Wu F. Y.and Gao S. (in press) Comparative Sr–Nd–Pb–Hf–Os isotopicsystematics of xenolithic peridotites from Yangyuan, NorthChina Craton. In: Goldschmidt Conference, Prague, Czech.
Liu R., Chen W., Sun J. and Li D. (1990) The K–Ar age andtectonic environment of Cenozoic volcanic rock in China. InThe Age and Geochemistry of Cenozoic Volcanic Rock in China
(ed. R. Liu). Seismology Published House, Beijing, pp. 1–43.
Lorand J. P. and Alard O. (2001) Platinum-group elementabundances in the upper mantle: new constraints from in situand whole-rock analyses of Massif Central xenoliths (France).Geochim. Cosmochim. Acta 65, 2789–2806.
Lorand J. P., Alard O. and Luguet A. (2010) Platinum-groupelement micronuggets and refertilization process in Lherzorogenic peridotite (northeastern Pyrenees, France). Earth
Planet. Sci. Lett. 289, 298–310.
Lorand J. P., Delpech G., Gregoire M., Moine B., O’Reilly S. Y.and Cottin J. Y. (2004) Platinum-group elements and themultistage metasomatic history of Kerguelen lithosphericmantle (South Indian Ocean). Chem. Geol. 208, 195–215.
Lorand J. P., Luguet A. and Alard O. (2008) Platinum-groupelements: a new set of key tracers for the Earth’s interior.Elements 4, 247–252.
Lu L. Z., Xu X. C. and Liu F. L. (1996) Early Precambrian
Khondalite Series of North China. Changchun PublishingHouse, Changchun, pp. 1–272.
Luguet A., Shirey S. B., Lorand J. P., Horan M. F. and Carlson R.W. (2007) Residual platinum-group minerals from highlydepleted harzburgites of the Lherz massif (France) and theirrole in HSE fractionation of the mantle. Geochim. Cosmochim.
Acta 71, 3082–3097.
McDonough W. F. and Sun S. S. (1995) The composition of theEarth. Chem. Geol. 120, 223–253.
Meibom A., Sleep N. H., Chamberlain C. P., Coleman R. G., FreiR., Hren M. T. and Wooden J. L. (2002) Re–Os isotopicevidence for long-lived heterogeneity and equilibration pro-cesses in the Earth’s upper mantle. Nature 419, 705–708.
Meisel T., Reisberg L., Moser J., Carignan J., Melcher F. andBrugmann G. (2003) Re–Os systematics of UB-N, aserpentinized peridotite reference material. Chem. Geol. 201,
161–179.
Meisel T., Walker R. J., Irving A. J. and Lorand J. P. (2001)Osmium isotopic compositions of mantle xenoliths: a globalperspective. Geochim. Cosmochim. Acta 65, 1311–1323.
Menzies M., Xu Y. G., Zhang H. F. and Fan W. M. (2007)Integration of geology, geophysics and geochemistry: a key tounderstanding the North China Craton. Lithos 96, 1–21.
Menzies M. A., Fan W.-M. and Zhang M. (1993) Paleozoic andCenozoic lithoprobes and the loss of >120 km of Archeanlithosphere, Sino-Korean craton, China. In Magmatic Processes
and Plate Tectonics (eds. H. M. Prichard, H. M. Alabaster, T.Harris and C. R. Neary). Geological Society, London, pp. 71–
Mapping lithospheric boundaries via Os isotopes 3901
O’Neill H. S. C. (1981) The transition between spinel lherzolite andgarnet lherzolite, and its use as a Geobarometer. Contrib.
Mineral. Petrol. 77, 185–194.
Parkinson I. J., Hawkesworth C. J. and Cohen A. S. (1998) Ancientmantle in a modern arc: osmium isotopes in Izu–Bonin–Mariana forearc peridotites. Science 281, 2011–2013.
Pearson D. G., Carlson R. W., Shirey S. B., Boyd F. R. and NixonP. H. (1995) Stabilization of Archean lithospheric mantle – aRe–Os isotope study of peridotite xenoliths from the KaapvaalCraton. Earth Planet. Sci. Lett. 134, 341–357.
Pearson D. G., Canil D. and Shirey S. B. (2003) Mantle samplesincluded in volcanic rocks: xenoliths and diamonds. In Treatise
on Geochemistry, The Mantle and Core, vol. 2 (ed. R. W.Carlson), pp. 171–276.
Pearson D. G., Irvine G. J., Ionov D. A., Boyd F. R. and DreibusG. E. (2004) Re–Os isotope systematics and platinum groupelement fractionation during mantle melt extraction: a study ofmassif and xenolith peridotite suites. Chem. Geol. 208, 29–59.
Pearson D. G. and Wittig N. (2008) Formation of Archaeancontinental lithosphere and its diamonds: the root of theproblem. J. Geol. Soc. 165, 895–914.
Puchtel I. and Humayun M. (2000) Platinum group elements inKostomuksha komatiites and basalts: implications for oceaniccrust recycling and core–mantle interaction. Geochim. Cosmo-
chim. Acta 64, 4227–4242.
Puchtel I. S., Walker R. J., James O. B. and Kring D. A. (2008)Osmium isotope and highly siderophile element systematics oflunar impact melt breccias: implications for the late accretionhistory of the Moon and Earth. Geochim. Cosmochim. Acta 72,
3022–3042.
Rehkamper M., Halliday A. N., Fitton J. G., Lee D. C., WienekeM. and Arndt N. T. (1999) Ir, Ru, Pt, and Pd in basalts andkomatiites: new constraints for the geochemical behavior of theplatinum-group elements in the mantle. Geochim. Cosmochim.
Acta 63, 3915–3934.
Reisberg L. and Lorand J. P. (1995) Longevity of sub-continentalmantle lithosphere from osmium isotope systematics in oro-genic peridotite massifs. Nature 376, 159–162.
Reisberg L., Zhi X. C., Lorand J. P., Wagner C., Peng Z. C. andZimmermann C. (2005) Re–Os and S systematics of spinelperidotite xenoliths from east central China: evidence forcontrasting effects of melt percolation. Earth Planet. Sci. Lett.
239, 286–308.
Rogers J. J. W. and Santosh M. (2003) Supercontinents in earthhistory. Gondwana Res. 6, 357–368.
Rudnick R. L., Shan G., Ling W. L., Liu Y. S. and McDonoughW. F. (2004) Petrology and geochemistry of spinel peridotitexenoliths from Hannuoba and Qixia, North China craton.Lithos 77, 609–637.
Rudnick R. L. and Walker R. J. (2009) Interpreting ages from Re–Os isotopes in peridotites. Lithos 112S, 1083–1095.
Santosh M., Sajeev K. and Li J. H. (2006) Extreme crustalmetamorphism during Columbia supercontinent assembly:evidence from North China Craton. Gondwana Res. 10, 256–
266.
Santosh M., Wilde S. A. and Li J. H. (2007) Timing ofPaleoproterozoic ultrahigh-temperature metamorphism in theNorth China Craton: evidence from SHRIMP U–Pb zircongeochronology. Precambrian Res. 159, 178–196.
Shirey S. B. and Walker R. J. (1998) The Re–Os isotope system incosmochemistry and high-temperature geochemistry. Annu.
Rev. Earth Planet. Sci. 26, 423–500.
Song Y. and Frey F. A. (1989) Geochemistry of peridotite xenolithsin basalt from Hannuoba, Eastern China – implications forsubcontinental mantle heterogeneity. Geochim. Cosmochim.
Acta 53, 97–113.
Standish J. J., Hart S. R., Blusztajn J., Dick H. J. B. and Lee K. L.(2002) Abyssal peridotite osmium isotopic compositions fromCr-spinel. Geochem. Geophys. Geosyst. 3.
Tang Y. H., Zhang H. F., Ying J. F., Zhang J. and Liu X. M.(2008) Refertilization of ancient lithospheric mantle beneath thecentral North China Craton: evidence from petrology andgeochemistry of peridotite xenoliths. Lithos 101, 435–452.
Tang Y. J., Zhang H. F., Nakamura E., Moriguti T., Kobayashi K.and Ying J. F. (2007) Lithium isotopic systematics of peridotitexenoliths from Hannuoba, North China Craton: implicationsfor melt–rock interaction in the considerably thinned litho-spheric mantle. Geochim. Cosmochim. Acta 71, 4327–4341.
Tao W. and Shen Z. K. (2008) Heat flow distribution in Chinesecontinent and its adjacent areas. Prog. Nat. Sci. 18, 843–849.
Tian Y., Zhao D. P., Sun R. M. and Teng J. W. (2009) Seismicimaging of the crust and upper mantle beneath the North ChinaCraton. Phys. Earth. Planet. In. 172, 169–182.
Walker R. J., Shirey S. B., Hanson G. N., Rajamani V. and HoranM. F. (1989) Re–Os, Rb–Sr, and O isotopic systematics of theArchean Kolar Schist Belt, Karnataka, India. Geochim. Cos-
mochim. Acta 53, 3005–3013.
Walker R. J., Storey M., Kerr A. C., Tarney J. and Arndt N. T.(1999) Implications of Os-187 isotopic heterogeneities in amantle plume: evidence from Gorgona Island and Curacao.Geochim. Cosmochim. Acta 63, 713–728.
Wan Y. S., Song B., Liu D. Y., Wilde S. A., Wu J. S., Shi Y. R.,Yin X. Y. and Zhou H. Y. (2006) SHRIMP U–Pb zircongeochronology of Palaeoproterozoic metasedimentary rocks inthe North China Craton: evidence for a major Late Palaeo-proterozoic tectonothermal event. Precambrian Res. 149, 249–
271.
Wilde S. A., Zhao G. C. and Sun M. (2002) Development of theNorth China Craton during the late Archaean and its finalamalgamation at 1.8 Ga: some speculations on its positionwithin a global Palaeoproterozoic supercontinent. Gondwana
Res. 5, 85–94.
Wu F. Y., Walker R. J., Ren X. W., Sun D. Y. and Zhou X. H.(2003) Osmium isotopic constraints on the age of lithosphericmantle beneath northeastern China. Chem. Geol. 196, 107–129.
Wu F., Zhao G., Wilde S. A. and Sun D. (2005a) Nd isotopicconstraints on crustal formation in the North China Craton. J.
Asian Earth Sci. 24, 523–545.
Wu F. Y., Lin J. Q., Wilde S. A., Zhang X. O. and Yang J. H.(2005b) Nature and significance of the Early Cretaceous giantigneous event in eastern China. Earth Planet. Sci. Lett. 233,
103–119.
Wu F. Y., Walker R. J., Yang Y. H., Yuan H. L. and Yang J. H.(2006) The chemical-temporal evolution of lithospheric mantleunderlying the North China Craton. Geochim. Cosmochim.
Acta 70, 5013–5034.
Xu W. L., Yang D. B., Gao S., Pei F. P. and Yu Y. (2010)Geochemistry of peridotite xenoliths in Early Cretaceous high-Mg# diorites from the Central Orogenic Block of the NorthChina Craton: the nature of Mesozoic lithospheric mantle andconstraints on lithospheric thinning. Chem. Geol. 270, 257–273.
Xu Y. G. (2001) Thermo-tectonic destruction of the Archaeanlithospheric keel beneath the Sino-Korean Craton in China:evidence, timing and mechanism. Phys. Chem. Earth A: Solid
Earth Geodesy 26, 747–757.
Xu Y. G., Blusztajn J., Ma J. L., Suzuki K., Liu J. F. and Hart S.R. (2008a) Late Archean to early Proterozoic lithosphericmantle beneath the western North China craton: Sr–Nd–Osisotopes of peridotite xenoliths from Yangyuan and Fansi.Lithos 102, 25–42.
Xu X. S., Griffin W. L., O’Reilly S. Y., Pearson N. J., Geng H. Y.and Zheng J. P. (2008b) Re–Os isotopes of sulfides in mantle
Yang J. H., O’Reilly S., Walker R. J., Griffin W., Wu F. Y., ZhangM. and Pearson N. (2010) Diachronous decratonization of theSino-Korean craton: geochemistry of mantle xenoliths fromNorth Korea. Geology 38, 799–802.
Yin C. Q., Zhao G. C., Sun M., Xia X. P., Wei C. J., Zhou X. W.and Leung W. H. (2009) LA-ICP-MS U–Pb zircon ages of theQianlishan complex: constrains on the evolution of theKhondalite Belt in the Western Block of the North ChinaCraton. Precambrian Res. 174, 78–94.
Zhai M. G. and Liu W. J. (2003) Palaeoproterozoic tectonic historyof the North China craton: a review. Precambrian Res. 122,
183–199.
Zhang H. F. (2005) Transformation of lithospheric mantle throughperidotite–melt reaction: a case of Sino-Korean craton. Earth
Planet. Sci. Lett. 237, 768–780.
Zhang H. F., Goldstein S. L., Zhou X. H., Sun M. and Cai Y.(2009) Comprehensive refertilization of lithospheric mantlebeneath the North China Craton: further Os–Sr–Nd isotopicconstraints. J. Geol. Soc. 166, 249–259.
Zhang H. F., Goldstein S. L., Zhou X. H., Sun M., Zheng J. P. andCai Y. (2008) Evolution of subcontinental lithospheric mantlebeneath eastern China: Re–Os isotopic evidence from mantlexenoliths in Paleozoic kimberlites and Mesozoic basalts. Con-
trib. Mineral. Petrol. 155, 271–293.
Zhang H. F., Sun M., Zhou X. H., Fan W. M., Zhai M. G. and YinJ. F. (2002) Mesozoic lithosphere destruction beneath theNorth China Craton: evidence from major-, trace-element andSr–Nd–Pb isotope studies of Fangcheng basalts. Contrib.
Mineral. Petrol. 144, 241–253.
Zhang W. H. and Han B. F. (2006) K–Ar chronology andgeochemistry of Jining Cenozoic basalts, Inner Mongolia, andgeodynamic implications. Acta Petrol. Sin. 22, 1597–1607.
Zhao D. P., Maruyama S. and Omori S. (2007) Mantle dynamics ofWestern Pacific and East Asia: insight from seismic tomogra-phy and mineral physics. Gondwana Res. 11, 120–131.
Zhao G. C., Sun M., Wilde S. A. and Li S. Z. (2005) Late Archeanto Paleoproterozoic evolution of the North China Craton: keyissues revisited. Precambrian Res. 136, 177–202.
Zhao G. C., Wilde S. A., Cawood P. A. and Lu L. Z. (2000)Petrology and P–T path of the Fuping mafic granulites:implications for tectonic evolution of the central zone of theNorth China craton. J. Metamorph. Geol. 18, 375–391.
Zhao G. C., Wilde S. A., Cawood P. A. and Sun M. (2001)Archean blocks and their boundaries in the North ChinaCraton: lithological, geochemical, structural and P–T pathconstraints and tectonic evolution. Precambrian Res. 107, 45–
73.
Zhao G. C., Wilde S. A., Guo J. H., Cawood P. A., Sun M. and LiX. P. (2010) Single zircon grains record two Paleoproterozoiccollisional events in the North China Craton. Precambrian Res.
177, 266–276.
Zheng J. P. (1999) Mesozoic and Cenozoic Lithospheric Mantle
Replacement and Thinning in Eastern China. Press of ChinaUniversity of Geosciences, p. 126.
Zheng J. P., Griffin W. L., O’Reilly S. Y., Yu C. M., Zhang H. F.,Pearson N. and Zhang M. (2007) Mechanism and timing oflithospheric modification and replacement beneath the easternNorth China Craton: peridotitic xenoliths from the 100 MaFuxin basalts and a regional synthesis. Geochim. Cosmochim.
Acta 71, 5203–5225.
Zheng J. P., O’Reilly S. Y., Griffin W. L., Lu F. X., Zhang M. andPearson N. J. (2001) Relict refractory mantle beneath theeastern North China block: significance for lithosphere evolu-tion. Lithos 57, 43–66.
EA-1. Modal mineralogy and petrography of the peridotites from the North China Craton
EA-2. Olivine compositions of the peridotites collected from the seven localities using an Electron Probe Microanalysis (EPMA)
EA-3. Mineral compositions of peridotites measured by EPMA: olivine (EA-3a), orthopyroxene (EA-3b), clinopyroxene (EA-3c), and spinel (EA-3d)
EA-4. Whole rock major element analyses for peridotites from the seven localities determined using the XRF techniques
EA-5. Trace element analyses of blanks and reference materials (EA-5a), and whole rock peridotites from Fansi and Hebi plus two Yangyuan peridotites (EA-5b) measured by ICP-MS at Northwest University, China
Figure Captions
Fig. EX-1. Chondrite-normalized REE patterns of whole rock Hannuoba (a; data from Rudnick et al., 2004), Yangyuan (b; data from Y.G. Xu et al., 2008 and this study), Fansi (c), Hebi (d) and Fushan (e; data from Xu et al., 2010) peridotites. Chondrite values are from McDonough and Sun (1995)
Fig. EX-2. Primitive mantle-normalized trace element diagrams of whole rock peridotites. Data sources are the same as Fig. EX-1. Primitive mantle values are from Lyubetskaya and Korenaga, 2007
Fig. EX-3. 187Os/188Os versus Fo of olivine (a) and Cr# of spinel (b). Data sources in addition to this study are: Hannuoba (Gao et al., 2002; Rudnick et al., 2004) and Yangyuan (Liu et al., 2010)
Fig. EX-4. 187Os/188Os versue 1/Os of low-HSE Fushan peridotites
2
Appendix A: Petrography of peridotites from the North China Craton
In addition to the main text, we briefly review the petrology of each suite in order from north
to south. Peridotites from Hannuoba are mostly fresh and large (10-60 cm in diameter), with
good preservation of sulfides, both as inclusions in silicates, and interstitial phases at grain
boundaries (e.g., Rudnick et al., 2004; X.S. Xu et al., 2008; Liu et al., 2010). Petrology and
whole rock compositions of all Hannuoba samples, including the Re-Os and HSE data, are from
the studies of Gao et al. (2002), Rudnick et al. (2004), Becker et al (2006) and Liu et al. (2010).
Like Hannuoba, peridotites from Yangyuan are generally fresh and large, with maximum
diameters typically greater than 50 cm (Y.G. Xu et al., 2008; Liu et al., 2010). However,
Yangyuan peridotites are characterized by poor preservation of sulfides (Liu et al., 2010). All
Yangyuan samples described here were previously described in Y.G. Xu et al. (2008) or Liu et al.
(2010).
Peridotites from Datong are generally fresh, with poor preservation of sulfides. Xenoliths are
small, ranging from less than a centimeter to several centimeters across. Only five relatively
large samples were prepared for whole rock major element and HSE analyses. Data for two
additional samples, analyzed only for Re-Os isotopic systematics, are reported here.
Like Datong samples, peridotites from Jining are relatively small (<10 cm). In addition,
Jining samples are moderately to heavily altered along grain boundaries, with poor preservation
of sulfides. The small size of the samples prevented us from making sufficient powder for whole
rock trace element analysis.
Peridotites from Fansi are generally moderately altered and relatively small (<15 cm in
diameter), with poor preservation of sulfides. Most samples have protogranular to
porphyroclastic textures, but a few lherzolites show a corona texture, where spinel breaks down
3
into tiny grains at grain boundaries. Three additional samples, analyzed only for Re-Os isotopic
systematics, are reported here.
Peridotites from Hebi are small (only a few centimeters across; many of them were sampled
by a rock drill from the lavas), fresh, and generally coarse-grained, harzburgites with a few
dunites. In contrast to the study of Zheng et al. (2001), who described five spinel lherzolites with
olivines of low forsterite contents (i.e., Fo = molar Mg/(Mg+Fe2+) x 100 = 88.6-91.4) and fertile
calculated whole rock compositions, none of the Hebi peridotites collected here were fertile
lherzolites. Only a few sulfides are present as inclusions in silicates (Zheng et al., 2007; this
study).
Xenoliths from Fushan are dominated by refractory harzburgites and cpx-poor lherzolites
with a few chromite-bearing dunites. Hydrous minerals (e.g., phlogopite and amphibole), may
occur in harzburgite and cpx-poor lherzolites, and were interpreted as secondary phases after
original mantle partial melting (Xu et al., 2010). Spinel grains that are in contact with phlogopite
or amphibole commonly break down into small grains of spinel or chromite (Xu et al., 2010). All
Fushan sample powders are from the study of Xu et al. (2010).
Appendix B: Analytical methods and results for whole rock trace element compositions
Whole rock trace element compositions were determined using an inductively coupled
plasma mass spectrometry (ICP-MS; Agilent 7500a) after acid digestion of powders in Teflon
bombs at Northwest University, China. Four reference samples (AGV-2, BHOV-2, BCR-2 and
RGM-1) were analyzed during the course of these analyses. The precision and accuracy was
generally better than 10%, with a majority of elements better than 5% relative to the reference
values (Rudnick et al., 2004). Exceptions are for a few elements such as Cr and Sn in BCR-2 (up
to 26% difference) and Ni and Mo in RGM-1 (up to 31% difference).
4
The results of trace element concentrations of whole rock samples as well as reference
materials are provided in the Electronic Annex EA-5. The peridotites in this study show
significant variations in both absolute concentrations of trace elements and chondrite-normalized
rare earth element (REE) patterns. The Hannuoba peridotites are characterized by total REE
concentrations of 1.7-17.5 ppm, with an average of 5.1 ppm (Rudnick et al., 2004). Their REE
patterns range from light REE (LREE) enriched to LREE depleted, with relatively flat heavy
REE (HREE) (Fig. EX-1a and EX-2a). The Yangyuan peridotites display lower total REE (0.3-
4.7 ppm, with an average of 1.9 ppm), but show similar variation in REE patterns as the
Hannuoba peridotites (Fig. EX-1b). Both Hannuoba and Yangyuan peridotites generally show
somewhat chaotic patterns in the primitive mantle-normalized trace element diagrams (Fig. EX-
2a and b), which likely reflects the comparable influence of partial melting and metasomatism on
these rocks in lithophile elements.
By comparison, the Fansi peridotites exhibit somewhat higher total REE abundances (3.0-
61.3 ppm, with an average of 11.0 ppm) than Hannuoba and Yangyuan peridotites. The low- and
high-Fo Fansi samples show similar REE patterns with uniformly enriched LREE and flat HREE
(Fig. EX-1c). One sample (FS-36), however, has high HREE, low MREE and strong enrichment
in La and Ce; in this respect, it is similar to some of the Hannuoba peridotites (e.g., DMP-60; Fig.
EX-1a). The low-Fo Fansi peridotites display negative anomalies of high field strength elements
(HFSE), such as Zr, Hf, Nb and Ta, a typical feature of mantle partial melting (Norman, 1998),
while the high-Fo Fansi samples do not. This contrasting feature may reflect that refractory
peridotites are more easily affected by metasomatic processes than fertile samples in lithophile
elements. Both Hebi and Fushan peridotites are characterized by low total REE abundances (1.1-
11.9 ppm, an average of 3.8 ppm for Hebi, and 0.6-5.2 ppm, an average of 2.3 ppm for Fushan),
strong LREE enrichment and a characteristic concave upwards HREE pattern (Fig. EX-1 d and
e). Similar to the high-Fo Fansi peridotites, the Hebi peridotites generally do not show negative
5
anomalies of HFSE, while the Fushan peridotites show a variation from negative to positive
anomalies HFSE.
In the Hebi suite, dunite HB-24 has low Fo (89.3), high Fe2O3 (10.6 %) and low Al2O3
(0.96 %) and CaO (0.44%), which are likely indicative of Fe enrichment, and has the least
radiogenic 187Os/188Os in the Hebi suite (0.1125). This sample has the most abundant REE
among the Hebi suite, which were likely enriched in the same process as Fe. At this rate, the
negative Ce anomaly (Fig. EX-1 d) might reflect the presence of Ce4+ (Ionov et al., 1995),
recording an oxidized environment during Fe enrichment. Importantly, the Fe enrichment event
did not significantly change the Os isotopic composition of the whole rock. In addition, the only
three Hebi samples in the “low Fo” group of Zheng et al. (2001) plot off the trends defined by
the Fo of olivine and Cr# of spinel (Fig. EX-3); the calculated whole-rock Al contents of these
samples are in the range of the high-Fo group, although the calculated CaO appears to be higher.
These observations suggest that the Fo of olivines of these low-Fo samples are simply due to
recent disequilibrium exchange. Thus, the few low-Fo Hebi samples reported in Zheng et al.
(2001) may not necessarily represent Phanerozoic mantle, but rather Fe-enriched Archean mantle.
For all the peridotitic samples in this study, the siderophile element Ni remains relatively
constant in the whole rock, while the lithophile elements V and Cr are depleted. The depletion
extent of V and Cr is qualitatively proportional to the degree of partial melting (Fig. EX-2).
Copper, as a chalcophile element, behaves incompatibly (depletion relative to Ni) to compatibly
(no depletion and, even enrichment relative to Ni) in mantle peridotites (Fig. EX-2).
6
References:
Becker, H., Horan, M. F., Walker, R. J., Gao, S., Lorand, J. P., and Rudnick, R. L., 2006. Highly siderophile element composition of the Earth's primitive upper mantle: Constraints from new data on peridotite massifs and xenoliths. Geochimica et Cosmochimica Acta, v. 70, p. 4528-4550.
Gao, S., Rudnick, R. L., Carlson, R. W., McDonough, W. F., and Liu, Y. S., 2002. Re-Os evidence for replacement of ancient mantle lithosphere beneath the North China craton. Earth and Planetary Science Letters, v. 198, p. 307-322.
Ionov, D. A., Prikhodko, V. S., and Oreilly, S. Y., 1995. Peridotite Xenoliths in Alkali Basalts from the Sikhote-Alin, Southeastern Siberia, Russia - Trace-Element Signatures of Mantle beneath a Convergent Continental-Margin. Chemical Geology, v. 120, p. 275-294.
Liu, J. G., Rudnick, R. L., Walker, R. J., Gao, S., Wu, F. Y., and Piccoli, P. M., 2010. Processes controlling highly siderophile element fractionations in xenolithic peridotites and their influence on Os isotopes. Earth and Planetary Science Letters, v. 297, p. 287-297.
Lyubetskaya, T., and Korenaga, J., 2007. Chemical composition of Earth's primitive mantle and its variance: 1. Method and results. Journal of Geophysical Research-Solid Earth, v. 112, p. -.
Norman, M. D., 1998. Melting and metasomatism in the continental lithosphere: laser ablation ICPMS analysis of minerals in spinel lherzolites from eastern Australia. Contributions to Mineralogy and Petrology, v. 130, p. 240-255.
Rudnick, R. L., Shan, G., Ling, W. L., Liu, Y. S., and McDonough, W. F., 2004. Petrology and geochemistry of spinel peridotite xenoliths from Hannuoba and Qixia, North China craton. Lithos, v. 77, p. 609-637.
Xu, W. L., Yang, D. B., Gao, S., Pei, F. P., and Yu, Y., 2010. Geochemistry of peridotite xenoliths in Early Cretaceous high-Mg# diorites from the Central Orogenic Block of the North China Craton: The nature of Mesozoic lithospheric mantle and constraints on lithospheric thinning. Chemical Geology, v. 270, p. 257-273.
Xu, X. S., Griffin, W. L., O'Reilly, S. Y., Pearson, N. J., Geng, H. Y., and Zheng, J. P., 2008. Re-Os isotopes of sulfides in mantle xenoliths from eastern China: Progressive modification of lithospheric mantle. Lithos, v. 102, p. 43-64.
Xu, Y. G., Blusztajn, J., Ma, J. L., Suzuki, K., Liu, J. F., and Hart, S. R., 2008. Late Archean to early proterozoic lithospheric mantle beneath the western North China craton: Sr-Nd-Os isotopes of peridotite xenoliths from Yangyuan and Fansi. Lithos, v. 102, p. 25-42.
Zheng, J. P., Griffin, W. L., O'Reilly, S. Y., Yu, C. M., Zhang, H. F., Pearson, N., and Zhang, M., 2007. Mechanism and timing of lithospheric modification and replacement beneath the eastern North China Craton: Peridotitic xenoliths from the 100 Ma Fuxin basalts and a regional synthesis. Geochimica et Cosmochimica Acta, v. 71, p. 5203-5225.
Zheng, J. P., O'Reilly, S. Y., Griffin, W. L., Lu, F. X., Zhang, M., and Pearson, N. J., 2001. Relict refractory mantle beneath the eastern North China block: significance for lithosphere evolution. Lithos, v. 57, p. 43-66.
