Lithos 146-147 (2012) 202–217

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Slab–mantle interaction for thinning of cratonic lithospheric mantle in North China:Geochemical evidence from Cenozoic continental basalts in central Shandong

Zheng Xu, Zi-Fu Zhao ⁎, Yong-Fei ZhengCAS Key Laboratory of Crust-Mantle Materials and Environments, School of Earth and Space Sciences, University of Science and Technology of China, Hefei 230026, China

⁎ Corresponding author. Tel.: +86 551 3600093; fax:E-mail address: zfzhao@ustc.edu.cn (Z.-F. Zhao).

0024-4937/$ – see front matter © 2012 Elsevier B.V. Alldoi:10.1016/j.lithos.2012.05.019

a b s t r a c t

a r t i c l e i n f o

Article history:Received 20 January 2012Accepted 17 May 2012Available online 28 May 2012

Keywords:Continental basaltSlab–mantle interactionMelt–peridotite reactionOceanic subductionCraton destruction

The North China Craton is one of the Archean cratons on Earth, and it has been reactivated since the Mesozoicwith significant thinning of the subcontinental lithospheric mantle (SCLM). While geophysical observationsindicate the effect of the Pacific plate subduction on the westward thinning, the coherent records are alsoevident from geochemistry of Cenozoic continental basalts in North China. This is delineated by a combinedstudy of whole-rock major-trace elements and Sr–Nd–Pb isotopes, mineral O isotopes as well as phenocrystolivine elements in Cenozoic basalts from the southeastern part of the North China Craton. The resultssuggest that the origin of Cenozoic basalts has bearing on the SCLM thinning in North China. The basaltsare alkalic and can be divided into basanite and alkali basalt. They are all characterized by the OIB-like patternof trace element distribution with no Nb–Ta anomalies and negative Pb anomalies. Nevertheless, the basaniteexhibits higher contents of Na2O and most incompatible elements than the alkali basalt. In addition, thebasanite has initial Sr isotope ratios of 0.7032 to 0.7033 and εNd(t) values of 5.5 to 6.0, which are more homo-geneous and depleted than those of 0.7038 to 0.7045 and −2.9 to 2.6 for the alkali basalt. On the other hand,the basanite has initial 206Pb/204Pb ratios of 17.867 to 17.995, which are higher than those of 16.752 to 17.320for the alkali basalt. These observations indicate that the two subtypes of basalts were derived from differentcompositions of mantle sources. Some olivine and plagioclase phenocrysts have lower δ18O values than nor-mal mantle values, suggesting incorporation of the recycled crustal materials into their mantle sources. Allphenocryst olivines have relatively high Ni contents and Fe/Mn ratios but low Ca andMn contents, suggestingthat the mantle sources are dominated by pyroxenites in lithology. Thus, the slab–mantle interaction issuggested for formation of the fertile and juvenile SCLM domains above an oceanic subduction zone.Pyroxene-rich lithologies would be generated by reaction of the juvenile SCLM peridotite with felsic meltsfrom partial melting of the subducting oceanic crust with rutile breakdown. The slab–mantle interactionwould occur during theMesozoic subduction of the Pacific plate beneath the eastern edge of Eurasian continent,leading to thinning of the cratonic lithospheric mantle in North China. Partial melting of the non-peridotitelithologies in the Cenozoic brought about the basanite and alkali basalt, providing a snapshot of the slab–mantleinteraction during the thinning.

© 2012 Elsevier B.V. All rights reserved.

1. Introduction

The North China Craton (NCC) is one of the Archean cratons onEarth and it was reactivated in the Phanerozoic (Griffin et al., 1998;Menzies and Xu, 1998; Menzies et al., 1993; Xu, 2001), with signifi-cant destruction in the Mesozoic (Gao et al., 2009; Menzies et al.,2007; Zheng, 2009). It is well established that the thickness of conti-nental lithosphere in North China was reduced from over 200 km to60–80 km during the Mesozoic, in which the ancient, cold and refrac-tory subcontinental lithospheric mantle (SCLM) was replaced byjuvenile, hot and fertile SCLM (e.g., Fan et al., 2000; Gao et al., 2002;Wu et al., 2006; Zheng et al., 2005). According to studies of seismic

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tomography, the cratonic lithosphere is thinner in eastern NorthChina than in western North China (Chen, 2009; Huang and Zhao,2009; Tian et al., 2009; Xu et al., 2009). From west to east, the thick-ness of continental lithosphere varies from 200–250 km in the Ordosbasin to 100–150 km in the Taihang orogen to 60–100 km in thenorthern China platform and b60 km in the Shandong peninsula andBohai basin. This variation clearly indicates the effect of the Pacificplate subduction on the SCLM thinning in North China.

There have been hot debates on the mechanism of thinning the cra-tonic lithospheric mantle in North China. One school advocated verticaldelamination of the lower continental crust (e.g., Gao et al., 2004, 2009;Wu et al., 2003; Xu et al., 2006), whereas the other hypothesizedthermal-chemical erosion by the upwelling asthenospheric mantle(e.g., Xu, 2001; Zhang et al., 2002; Zheng et al., 2005, 2006). The twoviewpoints involve controversial interpretations of geochemical datafor Phanerozoic igneous rocks in the North China Craton. Based on

203Z. Xu et al. / Lithos 146-147 (2012) 202–217

the geophysical observations of lithospheric thickness across the NorthChina Craton, on the other hand, it is inferred that westward subduc-tion of the Pacific plate has played the first-order role in thinning theSCLM of North China (Zhang et al., 2009; Zhu and Zheng, 2009). Inthis regard, it is possible that both the mechanical delamination andthermo-chemical erosion would only serve as the second-ordermechanisms for the SCLM thinning (Zheng and Wu, 2009). If so, it isimportant to test the effect of the Pacific plate subduction on theSCLM thinning in North China.

The geological record of the SCLM thinning in North China is theoccurrence of voluminous mafic magmatic rocks in the Mesozoic toCenozoic, with contrasting features of enrichment and depletion inboth incompatible elements and radiogenic isotopes. Because maficmelts are petrologically derived from partial melting of ultramaficlithology, geochemical constraints on the origin of continental maficrocks are a key to understanding geodynamics of the SCLM thinning.With regard to this issue, many studies have focused onMesozoic maficrocks in North China (e.g., Gao et al., 2009; Xu et al., 2009; Zhang, 2009;and references therein), but less attention has been paid to Cenozoicbasalts in this region. It is known that the Mesozoic mafic rocks arecharacterized not only by the arc-like patterns of trace element distribu-tion (i.e. the relative enrichment of LILE, Pb and LREE but the relativedepletion of HFSE), but also by high initial Sr isotope ratios and negativeεNd(t) values (e.g., Menzies et al., 2007; Xu, 2001; Zhang et al., 2003).In contrast, the Cenozoic basalts are rich in sodium, and show theoceanic island basalts (OIB)-like patterns of trace element distribution(i.e. the relative enrichment of LILE and LREE but the relative depletionof Pb and the no depletion of HFSE) and relatively depleted Sr–Ndisotope compositions (e.g., Liu et al., 2008; Tang et al., 2007; Xu et al.,2005; Zou et al., 2000). These geochemical contrasts suggest remark-able differences in their mantle sources between the Mesozoic andCenozoic mafic rocks, recording a tectonic transition beneath the NCCin this period.

It is geochemically known that the intraplate basalts with OIB-liketrace element characteristics cannot be directly generated by partialmelting of the normal asthenospheric mantle (e.g., Niu and O'Hara,2008; Zhang et al., 2009); the mid-oceanic ridge basalts (MORB) arecommonly assumed to be derived from partial melting of the normalasthenospheric mantle (e.g., Salters and Stracke, 2004; Workman andHart, 2005). In other words, the intraplate basalts require a specificmantle source that is significantly enriched in LILE and LREE relativeto the MORB-type asthenospheric mantle but not depleted in HFSErelative to the arc-type wedge mantle (Wang et al., 2011; Zhanget al., 2009; Zheng, 2012). With respect to the origin of enriched com-ponents, there are different hypotheses such as the ancient SCLM(Tang et al., 2006; Xu et al., 2005), the delaminated lower continental

Fig. 1. Distribution of Cenozoic continental basalts (a) in the North China Craton (modified2007).

crust (Gao et al., 2002; Liu et al., 2008; Zeng et al., 2010, 2011), andthe subducting oceanic crust (Wang et al., 2011; Zhang et al., 2009).These hypotheses involve different types of enriched components inmantle sources, either such incompatible elements as LILE and LREEor radiogenic Sr–Nd isotopes. Sometimes they concern either fertileor refractory lithology in ultramafic mantle sources. In this context, itis important to take into account the geochemical features of majorand trace elements in addition to those of radiogenic isotopes in conti-nental mafic igneous rocks. The slab–mantle interaction above subduc-tion zones for the formation of fertile and enrichedmantle sources mayhold a key to resolve the controversies (Zheng, 2012).

Understanding the origin of Cenozoic continental basalts in the NCChas great bearing not only on the nature of theirmantle sources but alsoon coherent slab–mantle interaction and the SCLM thinning. This paperpresents a combined study ofwhole-rockmajor-trace elements and Sr–Nd–Pb isotopes, phenocryst mineral O isotopes and phenocryst olivineelements for the Cenozoic basalts from the southeastern NCC. The re-sults not only provide insights into petrogenesis of these continental ba-salts, but also place constraints on themechanism of the SCLM thinningin North China. Consequently, the Cenozoic continental basalts in theeastern edge of NCC provide a genetic link between the slab–mantle in-teraction and the SCLM thinning.

2. Geological setting and samples

Tectonically, the North China Craton is bounded by the earlyPaleozoic Qilian orogen in the west, the late Paleozoic to MesozoicCentral Asian orogen in the north, and the Mesozoic Qinling–Dabie–Sulu orogen in the south (Zhao et al., 2001). The oldest crustalremnants in the NCC are approximate 3.8 Ga (Liu et al., 1992). In itsPrecambrian history, the NCC is divided into the eastern terrane, thewestern terrane and the Trans-North China Orogen between thetwo terranes (Zhao et al., 2001). The two terranes have the markeddifference in Paleoproterozoic–Archean metamorphic history (Geet al., 2003; Zhao et al., 1998, 1999, 2001). While the western andeastern terranes underwent a major metamorphic event at about2.5 Ga with an anticlockwise P–T–t path associated with magmaunderplating (Zhao et al., 1998, 1999), the Trans-North China Orogenis a collisional zone between the two terranes and experienced amajor metamorphic event at about 1.85 Ga with clockwise P–T–t path(Zhao et al., 2000).

Cenozoic basalts in the NCC are mainly located in east-centralLiaoning, eastern Jilin, northwest Hebei, eastern and northern Shanxias well as north-central Shandong (Fig. 1a). These basalts primarilyerupted in the Miocene to Holocene, with a few in the Paleogene(Chen et al., 2007). Mantle and crustal xenoliths are common in

after Zou et al., 2000; Xu et al., 2005), and (b) in Shandong province (after Hu et al.,

Fig. 2. Whole-rock Ar/Ar dating for a Cenozoic basalt (sample 07CL09) in the Changle–Linqu area, North China. (a) Age spectra, and (b) inverse isochron.

204 Z. Xu et al. / Lithos 146-147 (2012) 202–217

these basalts. The Changle–Linqu area is located geographically incentral Shandong, and tectonically at the eastern border of the NCCand close to the Tanlu fault zone (Fig. 1b). This is the main area of out-cropping Cenozoic basalts in Shandong province (Hu et al., 2007).

The Cenozoic basalts in the Changle–Linqu area are underlain bycoal-bearing lacustrine sediments of the Paleogene Wutu formation(Zhang et al., 2006). Volcanic activity occurred mainly in three pe-riods, which are the Niushan period (~14.5 Ma), the Shanwang peri-od (~10.0 Ma) and the Yaoshan period (~4.3 Ma), respectively(Wang and Li, 1996). The basalts of the Niushan period are widelydistributed in the Changle–Linqu area as lava platforms or volcaniccones, whereas those of the Yaoshan period are limited on the topsof several volcanic cones (Dong et al., 1999). The rock types are main-ly alkali olivine basalt and basanite. Lherzolite xenoliths are commonin basalts, and some of them contain sapphire.

In this study, four samples (07CL01, 07CL03, 07CL04, 07CL05)were collected from the Linqu area, and six samples (07CL07,07CL08, 07CL09, 07CL10, 07CL11, 07CL13) were collected from theChangle area. Based on petrographic observations and alkali contents,the Changle–Linqu basalts can be divided into two subtypes, namelybasanite and alkali basalt. These basalts are porphyritic, and composedof olivine, plagioclase and pyroxene phenocrysts in a groundmass ofolivine, pyroxene and glass. For the alkali basalts, the phenocrystsinclude mainly olivine and plagioclase with minor pyroxene. Olivinephenocrysts are subhedral to anhedral, most of them show fractures.Plagioclase exhibits as tabular with small olivine and pyroxene grainsfilling between them. Polysynthetic twinning is common in theplagioclase. Pyroxene is subhedral and a bit smaller than olivine, andit is discrete in the rocks. Compared to the alkali basalt, the sizes ofthe phenocryst mineral grains in the basanite are generally small,especially for plagioclase. Matrix is composed of fine grains of olivineand pyroxene as well as glass. In sample 07CL08, almost all olivinesare completely altered to iddingsites or partially altered to iddingsitesleaving fresh core as relicts.

3. Analytical methods

Whole-rock major elements were measured on fused glass discsby X-ray fluorescence spectrometer (XRF) at Physical and ChemicalAnalytical Center in University of Science and Technology of China(USTC), Hefei. Analytical precision was better than ±5%, estimatedfrom repeated analyses of two standards (granitic gneiss GBW07121and plagioclase amphibolite GBW07122). Trace elements were mea-sured by Finnigan Element inductively coupled plasma mass spec-trometry (ICP-MS) at State Key Laboratory of Lithospheric Evolutionin Institute of Geology and Geophysics, Chinese Academy of Sciences(CAS), Beijing. Powders (~40 mg) were dissolved in HNO3+HF for5 days in Savillex Teflon screw-cap beakers at ~200 °C. Then the solu-tions were dried and diluted to 50 ml for analysis. Indiumwas used asinternal standard to correct instrument drift. The analytical precisionswere better than ±5% for most elements, estimated from repeatedanalyses of the Chinese national standard reference materials of graniteGSR1 and basalt GSR3.

Whole-rock 40Ar/39Ar dating was made at State Kay Laboratoryin Institute of Geology and Geophysics, CAS, Beijing. The samplewas crushed and sieved between 40 and 80 mesh fractions (380 to200 μm). After all visible crystals and impurities were removed byhand under a binocular microscope, the fresh matrix was washedwith acetone in an ultrasonic bath for 20 min. To remove possiblealteration, the matrix sample was washed with 5% HNO3 in an ultra-

Notes to Table 1:a Chinese national standard reference material. 07GB01 and 07GB02 are granitic gneiss

granite and basalt, respectively.b Mg number values were calculated assuming Fe2+/ΣFe=0.9.c (La/Yb)n and (La/Sm)n were calculated using the chondrite REE data of McDonough an

sonic bath for 20 min. The grains were then rinsed with distilledwater and dried. Approximately 3–4 mg of sample was wrapped inAl foil, and irradiated together with TCR-sanidine standards withassigned age of 28.34±0.28 Ma (Renne et al., 1998), and opticalCaF2 and K-glass monitors in position H8 of the 49-2 reactor, Beijing,China, for 30 h with 0.5 mm cadmium foil shield. Technical detailsof the step-heating analysis for the Ar/Ar dating were outlined inHe et al. (2008). 40Ar/39Ar step-heating analyses by furnace wereperformed on a MM5400 mass spectrometer operating in thestatic mode. The total system blanks (1000 °C, 20 min) were in therange of 4.9 to 5.8×10−16 mol for mass 40, 0.9 to 1.4×10−18 molfor mass 39, 8.7 to 9.2×10−19 mol for mass 37, and 1.8 to2.1×10−18 mol for mass 36. Mass discrimination (0.009934 to0.009958 per atomic mass unit) was monitored by analysis of 40Ar/36Ar air pipette aliquots each day. Ca, K correction factors were calcu-lated from the CaF2 and K-glass monitors: (40Ar/39Ar)K=8.8×10−4,(39Ar/37Ar)Ca=7.24×10−4, (36Ar/37Ar)Ca=2.39×10−4. The datawere corrected for system blanks, mass discriminations, interferingCa, K derived argon isotopes, and the decay of 37Ar since the time ofthe irradiation. The decay constant used throughout the calculationsis λ=5.543×10−10 a−1, as recommended by Steiger and Jaeger(1977). The uncertainty of J-value (±0.2% to ±0.5% in this work) isone standard deviation of mean; this was propagated into the finalplateau and isochron ages, and contributed about 40% to the totaluncertainty in these age determinations. The uncertainties of theages were reported as internal error which combines the analyticalerror and the error on the J-value. The plateau and isochron ageswere calculated using ArArCALC (Koppers, 2002).

GBW07121 and plagioclase amphibolite GBW07122, respectively. GSR1 and GSR3 are

d Sun (1995).

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Table 1Major and trace element compositions of Cenozoic basalts at Changle–Linqu in North China.

Rock type

Alkali basalt Basanite Standarda

Sample

07CL01 07CL07 07CL10 07CL11 07CL13 07CL03 07CL04 07CL05 07CL08 07CL09 07GB01 07GB02

Major element (wt.%)

SiO2 46.92 47.58 45.68 45.79 44.94 42.40 42.91 43.04 42.91 42.14 66.22 49.51 TiO2 2.20 1.89 2.23 2.52 2.17 2.55 2.56 2.75 2.89 2.70 0.29 0.74 Al2O3 14.76 13.91 13.29 13.30 12.84 13.66 13.83 13.50 13.38 13.60 16.43 13.64 Fe2O3

T

11.98 12.17 12.93 13.01 12.93 14.12 14.01 12.92 14.23 13.19 3.17 14.30 MnO 0.17 0.17 0.17 0.18 0.17 0.25 0.24 0.19 0.20 0.22 0.06 0.21 MgO 8.35 9.27 10.91 9.21 11.82 7.79 7.75 11.09 9.69 9.03 1.77 7.25 CaO 8.91 8.08 8.18 8.23 8.15 7.83 8.17 8.79 9.04 8.66 2.50 9.02 Na2O 3.25 2.91 2.44 2.61 2.22 5.43 4.82 3.19 4.75 4.76 5.20 1.92 K2O 1.06 1.24 1.49 1.48 1.47 2.90 2.97 1.82 0.88 2.19 2.57 0.45 P2O5 0.41 0.37 0.37 0.42 0.35 1.10 1.12 0.58 0.73 1.07 0.11 0.08 LOI 1.84 2.23 2.12 2.95 2.61 1.68 1.75 1.92 1.47 2.09 1.28 1.06 Total 99.85 99.82 99.81 99.69 99.67 99.71 100.14 99.79 100.16 99.65 99.60 98.19 K2O/Na2O 0.33 0.43 0.61 0.57 0.66 0.53 0.62 0.57 0.19 0.46 Mg#b 61 63 65 61 67 55 55 65 60 60 Fe/Mn 64.79 65.82 67.89 67.14 67.89 51.84 51.86 62.07 65.91 53.66

Rock type

Alkali basalt Basanite Standarda

Sample

07CL01 07CL07 07CL10 07CL11 07CL13 07CL03 07CL04 07CL05 07CL08 07CL09 GSR1 GSR3

Trace element (ppm)

Li 6.24 7.24 7.48 6.55 5.77 9.75 7.43 5.23 6.88 10.09 130.58 9.38 Be 1.55 1.26 1.33 1.58 1.18 4.21 4.38 2.05 2.64 3.78 12.40 2.49 Sc 23.48 22.69 23.65 22.12 20.70 14.10 14.10 24.04 21.06 17.13 6.23 15.65 V 208.59 197.12 212.07 221.03 199.63 169.97 162.29 202.34 234.28 182.56 21.84 170.74 Cr 294.22 270.34 351.16 275.63 359.90 157.76 154.18 294.70 230.01 185.58 2.99 141.86 Co 45.12 49.00 55.64 48.89 54.76 41.30 40.43 47.73 52.98 42.38 3.15 45.26 Ni 141.64 181.73 236.07 181.28 240.18 125.19 119.98 208.90 207.38 134.93 1.02 135.24 Cu 71.01 60.81 68.11 73.36 58.66 44.88 40.81 54.04 53.34 45.03 2.53 49.41 Zn 97.00 100.14 100.24 108.09 92.34 128.15 124.84 88.63 115.04 113.03 26.50 147.65 Ga 20.82 19.07 19.68 20.72 17.52 23.85 24.33 19.55 23.47 23.12 19.85 24.45 Rb 19.17 17.15 16.30 15.97 15.84 36.94 37.33 21.72 10.97 30.23 473.49 39.18 Sr 541.63 513.28 474.92 461.16 465.19 1124.63 1241.13 674.63 893.08 1082.90 104.28 1060.95 Y 18.88 18.18 17.19 19.78 15.35 27.12 28.12 21.02 20.28 26.17 62.62 22.34 Zr 164.53 140.09 146.80 161.68 130.00 391.23 419.76 206.22 254.82 354.02 170.22 275.19 Nb 29.65 24.19 27.91 31.59 24.69 90.48 89.78 46.88 59.42 88.11 38.75 67.12 Cs 0.17 0.16 0.20 0.17 0.16 0.48 0.51 0.24 0.34 0.52 40.67 0.48 Ba 242.95 297.92 256.74 257.51 232.85 372.44 420.88 335.82 314.58 447.67 326.24 515.55 Hf 3.86 3.49 3.85 4.21 3.25 8.62 9.33 4.84 5.96 7.88 6.17 6.55 Ta 1.90 1.57 1.86 2.12 1.66 6.10 6.46 3.18 3.98 6.04 6.85 4.22 Tl 0.07 0.08 0.07 0.07 0.07 0.08 0.09 0.06 0.06 0.08 1.86 0.06 Pb 2.63 3.19 2.61 2.57 2.30 5.19 5.89 2.97 3.19 5.49 30.35 4.51 Bi 0.02 0.02 0.01 0.01 0.02 0.02 0.02 0.01 0.01 0.02 0.51 0.02 Th 2.23 2.32 2.17 2.51 1.96 7.63 9.55 4.27 4.53 9.20 50.20 5.89 U 0.69 0.58 0.67 0.74 0.59 2.47 2.93 1.21 1.45 2.63 17.59 1.37

REE (ppm)

La 23.20 22.94 21.23 23.62 18.71 70.73 71.77 35.70 42.89 67.60 52.65 56.05 Ce 48.05 47.06 43.90 48.54 38.91 130.64 139.26 66.70 83.06 124.52 102.94 105.03 Pr 5.68 5.71 5.31 5.92 4.75 14.39 15.58 8.05 9.83 14.62 12.49 12.76 Nd 22.86 22.46 21.79 24.42 18.84 53.40 56.58 31.22 37.15 54.34 45.45 51.25 Sm 5.13 5.12 4.94 5.53 4.39 9.88 11.11 6.61 7.68 10.76 9.40 9.57 Eu 1.68 1.65 1.60 1.80 1.42 3.12 3.31 2.06 2.37 3.12 0.80 3.16 Gd 4.85 4.79 4.65 5.30 4.21 8.78 9.47 6.01 6.86 8.86 8.73 8.71 Tb 0.73 0.70 0.69 0.80 0.62 1.18 1.29 0.85 0.93 1.18 1.59 1.15 Dy 3.95 3.87 3.64 4.45 3.42 5.92 6.50 4.53 4.73 5.94 10.06 5.48 Ho 0.71 0.71 0.65 0.83 0.63 1.01 1.13 0.83 0.79 1.01 2.09 0.88 Er 1.72 1.74 1.59 1.98 1.56 2.42 2.60 1.99 1.83 2.38 6.21 1.97 Tm 0.23 0.24 0.21 0.26 0.21 0.29 0.34 0.27 0.22 0.30 1.02 0.24 Yb 1.29 1.42 1.20 1.56 1.25 1.52 1.83 1.60 1.18 1.68 6.93 1.25 Lu 0.17 0.20 0.17 0.21 0.17 0.20 0.24 0.22 0.16 0.23 1.08 0.16 (La/Yb)nc 12.20 10.94 11.98 10.32 10.18 31.63 26.65 15.19 24.65 27.27 (La/Sm)nc 2.82 2.80 2.68 2.67 2.66 4.47 4.03 3.37 3.49 3.92 La/Nb 0.78 0.95 0.76 0.75 0.76 0.78 0.80 0.76 0.72 0.77 Nb/Ta 15.65 15.46 15.04 14.89 14.85 14.82 13.89 14.72 14.95 14.59 Nb/U 42.72 41.64 41.77 42.58 41.78 36.57 30.63 38.64 41.07 33.50 Ce/Pb 18.26 14.73 16.84 18.90 16.91 25.18 23.63 22.43 26.05 22.67 Sr/Ce 11.27 10.91 10.82 9.50 11.96 8.61 8.91 10.11 10.75 8.70 Rb/Sr 0.04 0.03 0.03 0.03 0.03 0.03 0.03 0.03 0.01 0.03 Sm/Nd 0.22 0.23 0.23 0.23 0.23 0.18 0.20 0.21 0.21 0.20 U/Pb 0.26 0.18 0.26 0.29 0.26 0.48 0.50 0.41 0.45 0.48 Th/Pb 0.85 0.73 0.83 0.98 0.85 1.47 1.62 1.43 1.42 1.67

Fig. 3. Total alkalis vs. SiO2 diagram for Changle–Linqu basalts in North China. Theframework is after Le Maitre (2002).

206 Z. Xu et al. / Lithos 146-147 (2012) 202–217

Whole-rock Sr–Nd–Pb isotope analyses were chemically made atCAS Key Laboratory of Crust–Mantle Materials and Environments inUSTC, Hefei. Sample powders (~100 mg) were dissolved in HClO4+HFfor ~7 days in Savillex Teflon screw-cap beakers at ~120 °C. Then thesolutions were dried and dissolved in HCl for the separation of traceelements. Sr was carried out from solutions in turn by cation-exchange column filled by Bio-Rad resin. Then the leached solutionwas poured into HDEHP resin to leach Nd. Another 100 mg powderwas used to separate Pb, which was leached from the solutions thenpurified twice through AG1-X8 resin. Sr and Pb isotope ratios were de-termined by Finnigan MAT-262 thermal ionization mass spectrometry(TIMS) at CAS Key Laboratory of Crust-Mantle Materials and Environ-ments in USTC, Hefei. A 86Sr/88Sr ratio of 0.1194was used for normaliza-tion. Themeasured 87Sr/86Sr ratio of standard NBS987 and the precisionof measurements are 0.710298±20 and ~0.002%, respectively. For Pbisotopes, measured 206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb ratios ofstandard NBS981 are 16.917±0.02, 15.466±0.02 and 36.626±0.04,respectively. The precision of measurements is better than ±0.025%(mostly better than±0.015%) for all three Pb isotope ratios. The Nd iso-tope ratios were measured on a mass spectrometer Finnigan TRITON atTianjin Institute of Geology and Mineral Resources in Chinese Ministryof Land and Resources, Tianjin. A 146Nd/144Nd ratio of 0.7219 was usedfor normalization. The measured 143Nd/144Nd ratio of standard La Jollaand precision of measurement are 0.511851±45 and 0.0001% to0.0003%, respectively.

Mineral O isotope ratios were determined by the laser fluorinationmethod at CAS Key Laboratory of Crust–Mantle Materials and Envi-ronments in USTC, Hefei. Minerals with weights of 1.5 to 2.0 mgwere reacted with BrF5 in vacuum, and extracted O2 was directlytransferred to a mass spectrometer for the measurement of O isotoperatios (Zheng et al., 2002). A Finnigan Delta XP mass spectrometrywas used in the measurement of O isotope ratios. The results arereported in the δ18O notation relative to VSMOW. Analytical errorsare better than ±0.1‰. Reference minerals used in the laboratoryare as follows: garnet UWG‐2 with δ18O=5.8‰ (Valley et al., 1995)and δ18O=3.6‰ for garnet 04BXL07 (Gong et al., 2007).

Olivine elements were analyzed in suit by JXA-8100 electronicprobe at State Key Laboratory of Geologic Processes and MineralResources in China University of Geosciences, Wuhan. Acceleratevoltage and probe diameter are 15 kV and 3 μm, respectively. Analyticalerror is better than ±5%.

4. Results

4.1. Ar/Ar dating

The results of 40Ar/39Ar analyses for the basalt are listed in Table S1and plotted as the age spectrum and isochron diagram in Fig. 2.The glass matrix separated from lava 07CL09 gave a concordant agespectrum (Fig. 2). Nine consecutive steps which account for 87.8% ofthe total 39Ar released, define a plateau age of 19.0±0.6 Ma withMSWD=0.28 (Fig. 2a). An inverse isochron age of 18.6±1.7 Mawith MSWD=0.31 is calculated from the plateau steps (Fig. 2b), ingood agreement with the plateau age. An 40Ar/36Ar intercept value of296.5±3.8 is obtained, which is consistent with the atmospheric Arisotope ratio. Thus, there is no contamination of apparent excessargon in the sample. Because no assumption was made on the initial40Ar/36Ar ratio or its uncertainty, we interpret the isochron age 18.6±1.7 Ma as the best estimate for the time since eruption of the lava.

4.2. Major and trace elements

Major element data for the Changle–Linqu basalts are listed inTable 1. In terms of the nomenclature of Le Maitre (2002), theybelong to tephrite basanite and alkali basalt (Fig. 3). The basaltshave high alkali contents and lie above the line separating alkali and

tholeiitic basalts. They exhibit low K2O/Na2O ratios of 0.33 to 0.66for the alkali basalt and 0.19 to 0.62 for the basanite. The basanitehas Na2O contents of 3.19 to 5.43‰, which are systematically higherthan those of 2.22 to 3.25% for the alkali basalt. The ranges of MgOcontents are similar between the basanite (7.75 to 11.09%) and alkalibasalt (8.35 to 11.82%). Compared to the alkali basalt at the sameMgO,the basanite tends to exhibit higher contents of TiO2, Fe2O3

T, MnO, Na2Oand P2O5 but lower SiO2 contents, Mg# and Fe/Mn ratios (Table 1 andFig. 4). There are mildly negative correlations between Al2O3, Na2O,P2O5 and MgO for both basanite and alkali basalt (Fig. 4), but no sig-nificant correlations between the other oxides and MgO.

Trace element data for the Changle–Linqu basalts are listed inTable 1. Both basanite and alkali basalt exhibit relative enrichmentof LREE (light rare earth elements) without obvious Eu anomalies inthe chondrite-normalized REE diagram (Fig. 5a), which are similarto common OIB. While the two types of basalts show similar rangesin contents of HREE (heavy rare earth elements), the basanite hashigher LREE contents than the alkali basalt. As a consequence, the for-mer has higher (La/Yb)N ratios of 15.2 to 31.6 than those of 10.2 to12.2 for the latter. In the primitive mantle-normalized spidergram,all the basalts exhibit different degrees of enrichment in LILE (largeion lithosphile elements) and LREE relative to HREE, and displayOIB-like patterns with no Nb–Ta depletion but negative Pb anomalies(Fig. 5b). For most incompatible elements, the basanite generallyhas higher contents than the alkali basalt. All the basalts have Nb/Taratios of 13.89 to 15.65, which are lower than a common ratio of17.8 for the primitive mantle (McDonough and Sun, 1995). Exceptfor two basanite samples 07CL04 and 07CL09 with low Nb/U ratiosof 30.63 and 33.50, the other samples have high Nb/U ratios of 36.57to 42.72 (Table 1), which are close to the field of OIB with Nb/U ratiosof 47±10 (Hofmann et al., 1986).

4.3. Whole-rock Sr–Nd–Pb isotopes

Whole-rock Rb–Sr and Sm–Nd isotope data for the Changle–Linqubasalts are listed in Table 2. Initial Sr and Nd isotope ratios were cal-culated at t=20 Ma. The basanite has homogeneous initial 87Sr/86Srratios of 0.7032 to 0.7033 and εNd(t) values of 5.5 to 6.0. In contrast,the alkali basalt has relatively elevated initial 87Sr/86Sr ratios of0.7038 to 0.7045 but decreased εNd(t) values of −2.9 to 2.6,suggesting that the basanite is more depleted in the Sr–Nd isotopesthan the alkali basalt. Taking the both types together, there is a nega-tive correlation between the initial Sr and Nd isotope ratios (Fig. 6a),falling within the known ranges for Cenozoic basalts in the NCC.

Whole-rock Pb isotope data for the Changle–Linqu basalts arelisted in Table 3. Initial Pb isotope compositions were also calculatedat t=20 Ma. The basanite has 206Pb/204Pb ratios of 17.867 to 17.995,

Fig. 4. Major element oxides vs. MgO diagram for Changle–Linqu basalts in North China.

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207Pb/204Pb ratios of 15.451 to 15.503 and 208Pb/204Pb ratios of37.744 to 37.927. The alkali basalt has 206Pb/204Pb ratios of 16.752to 17.320, 207Pb/204Pb ratios of 15.299 to 15.550 and 208Pb/204Pb ra-tios of 37.312 to 38.039. It appears that the basanite is more enrichedin radiogenic Pb isotopes than the alkali basalt. The Pb isotope ratiosfor the Changle–Linqu basalts also fall within the known ranges forother Cenozoic basalts in the NCC. In the 207Pb/204Pb vs. 206Pb/204Pbdiagram (Fig. 6b), most samples fall close to or slightly above theNorthern Hemisphere Reference Line (NHRL) as established by Hart(1984). In the 208Pb/204Pb vs. 206Pb/204Pb diagram (Fig. 6c), all sam-ples fall above the NHRL.

4.4. Phenocryst O isotopes

Table 4 lists O isotope data for olivine and plagioclase phenocrystsfrom the Changle–Linqu basalts. The basanite has olivine and plagio-clase δ18O values of 4.15 to 5.70‰ and 5.15 to 6.07‰, respectively.The alkali basalt has olivine and plagioclase δ18O values of 4.09to 5.28‰ and 5.31 to 6.55‰, respectively. For olivine, δ18O valuesof 5.2±0.2‰ are considered the best estimate for the O isotope com-position of normal mantle reservoirs (Eiler et al., 1996). Using a frac-tionation factor of 0.75‰ between plagioclase and olivine at thetemperature 1000 to 1200 °C (Macpherson et al., 1998; Thirlwall etal., 1997; Zheng, 1993), the plagioclase δ18O values for normal mantlereservoirs would be 5.95±0.20‰. In this regard, most samples have

olivine and plagioclase δ18O values lower than those for the normalmantle, and the others are either similar to or higher than the normalmantle values (Fig. 7a, b). In the δ18O–δ18O diagram, there is a nega-tive correlation between the δ18O values of olivine and plagioclasefor the alkali basalt (Fig. 7c). Some samples yield unreasonably highO isotope apparent temperatures between olivine and plagioclase,suggesting O isotope disequilibrium between them.

4.5. Phenocryst olivine chemistry

Olivine from the alkali basalt and basanite has Fo values of 55 to83 and 71 to 88, respectively (Table S2). For single samples, low Fovalues are always obtained on relatively smaller grains, especiallyfor the basanite. Compared to olivine from the alkali basalt, olivinefrom the basanite has lower Ni contents and Fe/Mn ratios but higherMn contents at the same Fo values, especially when the Fo values arelower than 80 (Fig. 8a, b, d). The Ca contents of olivine from thebasanite are comparable with, or lower than, olivine from the alkalibasalt at the same Fo values when the Fo values are lower and higherthan 80, respectively (Fig. 8c).

5. Discussion

Geochemical studies of mafic magmatic rocks can provide valuableinformation on the nature and origin of their mantle sources. Variousmodels were proposed for interpretation of their petrogenesis by

Fig. 5. Spidergrams of trace element distribution for Changle–Linqu basalts in NorthChina. (a) Chondrite-normalized REE patterns, with the chondrite REE compositionis after Sun and McDonough (1989). (b) Primitive mantle-normalized trace elementpatterns, with the primitive mantle composition are after McDonough and Sun (1995).OIB data are from Sun and McDonough (1989). Melt-extracted lower crust modalis after Liu et al. (2008). Partition coefficients between garnet and basaltic melt areafter Adam and Green (2006) whereas partition coefficients between garnet and felsicmelt and partition coefficients of clinopyroxene, orthopyroxene and rutile are after Bédard(2006).

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envisaging different types of crust-mantle interaction (e.g., DePaolo,1981; Gray et al., 1981; Ringwood, 1990; Taylor, 1980; Zheng, 2012;Zindler et al., 1981). However, most of the models were principallyinferred from, or tested using, radiogenic isotopic data alone formafic rocks. For example, neither major elements nor trace elements

Table 2Rb–Sr and Sm–Nd isotope ratios of Cenozoic basalts at Changle–Linqu in North China.

Rock type Alkali basalt

Sample 07CL01 07CL07 07CL10 07CL11 07CL13

Rb (ppm) 19.17 17.15 16.30 15.97 15.84Sr (ppm) 541.63 513.28 474.92 461.16 465.1987Rb/86Sr 0.2685 0.2535 0.2604 0.2627 0.258(87Sr/86Sr)m 0.704062 0.704590 0.704078 0.703908 0.7042σ 0.000015 0.000014 0.000015 0.000014 0.000(87Sr/86Sr)ta 0.703986 0.704519 0.704005 0.703835 0.704Sm (ppm) 5.13 5.12 4.94 5.53 4.39Nd (ppm) 22.86 22.46 21.79 24.42 18.84147Sm/144Nd 0.1386 0.1408 0.1401 0.1399 0.143(143Nd/144Nd)m 0.512695 0.512480 0.512701 0.512765 0.5122σ 0.000003 0.000002 0.000001 0.000002 0.000(143Nd/144Nd)ta 0.512691 0.512476 0.512697 0.512761 0.512εNd(t) 1.26 −2.94 1.36 2.62 1.34TDM (Ma) 923 1398 930 795 982

a Initial isotope compositions were calculated using t=20 Ma.

were quantitatively taken into account when applying the model ofcrustal contamination to the interpretation of radiogenic isotopevariations. In this regard, to reasonably decipher the origin of maficrocks requires an integrated study of major and trace elements aswell as stable and radiogenic isotopes (e.g., Wang et al., 2011; Zhanget al., 2009).

Similar to oceanic island basalts, continental basalts are also char-acterized by larger variations in both trace element and radiogenicisotope compositions (Farmer, 2003; Hofmann, 1997). Nevertheless,a prominent feature is that continental basalts commonly show theOIB-like patterns of trace element distribution. As shown in Figs. 4to 7, the Cenozoic continental basalts at Changle–Linqu in easternNCC have considerable variations in whole-rock major-trace ele-ments and Sr–Nd–Pb isotopes as well as phenocryst O isotopes.These variations can be ascribed to heterogeneity in mantle sourcecompositions, variations in P–T conditions of partial melting, andmagmatic processes such as crystal fractionation and crustal contam-ination. The relative contribution of these factors to the basalts is dis-cussed separately below. As shown in Table 2, all the Changle–Linqubasalts have relatively low LOI values of 1.47 to 2.95% (Table 2),suggesting that they have experienced insignificant postmagmaticalteration. Although olivine phenocrysts in sample 07CL08 exhibitvarying degrees of alteration to the low-T iddingsite, the other samplesshow no obvious postmagmatic alteration. Therefore, the geochemicalcompositions of basalts in question can be used to decipher theirmagma sources

5.1. Continental crustal contamination

Many studies have revealed that Cenozoic basalts in eastern Chinawere not significantly affected by crustal contamination duringmagma ascent (e.g., Tang et al., 2006; Wang et al., 2011; Zhanget al., 2009; Zhi et al., 1990; Zou et al., 2000). The occurrence of man-tle xenoliths in the Changle–Linqu basalts suggests that the hostmagmas would ascend fast and did not suffer significant contamina-tion by the continental crust. The basalts exhibit depletion of suchLILE as K, Rb and Cs relative to the other incompatible elements inthe spidergram (Fig. 5b), suggesting insignificant contaminationfrom the upper continental crust. In addition, Nb/U and Ce/Pb ratioscan be used to evaluate the possibility of crustal contaminationbecause the upper continental crust has much lower Nb/U and Ce/Pb ratios than the upper mantle (Hofmann, 1988; Rudnick and Gao,2003; Salters and Stracke, 2004). Some basalt samples in this studyhave slightly lower Nb/U and Ce/Pb ratios than MORB and OIB(Fig. 9a, b). However, there are no negative correlations betweenNb/U, Ce/Pb and SiO2 for the basalts (Fig. 9c, d), indicating insignifi-cant crustal contamination during magma ascent. The alkali basalt

Basanite

07CL03 07CL04 07CL05 07CL08 07CL09

36.94 37.33 21.72 10.97 30.231124.63 1241.13 674.63 893.08 1082.90

4 0.2492 0.2281 0.2442 0.0932 0.2117255 0.703284 0.703331 0.703403 0.703348 0.703343015 0.000015 0.000015 0.000015 0.000014 0.000015183 0.703214 0.703267 0.703335 0.703322 0.703284

9.88 11.11 6.61 7.68 10.7653.40 56.58 31.22 37.15 54.34

8 0.1142 0.1212 0.1306 0.1276 0.1222700 0.512933 0.512923 0.512917 0.512911 0.512933002 0.000001 0.000003 0.000003 0.000003 0.000002696 0.512929 0.512919 0.512913 0.512907 0.512929

5.96 5.74 5.61 5.50 5.94334 376 428 424 362

Fig. 6. Sr–Nd–Pb isotope ratios for Changle–Linqu basalts in North China. NHRL is afterHart (1984). Data for Cenozoic basalts of the NCC are from database GEROC (http://georoc.mpch-mainz.gwdg.de/georoc/Entry.html). Pacific MORB data are from Whiteet al. (1987), and Indian MORB data are from Mahoney et al. (1989, 1992).

Table 3U–Th–Pb isotope compositions of Cenozoic basalts at Changle–Linqu in North China.

Alkali basalt

07CL01 07CL07 07CL10 07CL11

U (ppm) 0.69 0.58 0.67 0.74Th (ppm) 2.23 2.32 2.17 2.51Pb (ppm) 2.63 3.19 2.61 2.57238U/204Pb 47.010 31.898 45.222 51.046206Pb/204Pbm 17.325 16.851 17.325 17.3012σ (%) 0.014 0.013 0.015 0.015206Pb/204Pbta 17.179 16.752 17.184 17.143235U/204Pb 0.034 0.023 0.032 0.037207Pb/204Pbm 15.379 15.300 15.349 15.5512σ (%) 0.015 0.015 0.015 0.015207Pb/204Pbta 15.378 15.299 15.348 15.550232Th/204Pb 151.307 127.307 146.814 172.533208Pb/204Pbm 38.189 37.446 37.617 37.4832σ (%) 0.015 0.014 0.016 0.015208Pb/204Pbta 38.039 37.320 37.472 37.312

a Initial isotope compositions were calculated using t=20 Ma.

Table 4The oxygen isotope composition of phenocryst from Cenozoic basalts at Changle–Linquin North China.

Sample δ18OOl (‰) δ18OPl (‰)

Alkali basalt07CL01 4.09 6.5507CL07 4.96 6.0607CL10 4.60 5.9307CL11 5.28 5.3107CL13 4.56 5.61

Basanite07CL03 5.70 6.0707CL04 4.8407CL05 4.1907CL08 5.23 5.1507CL09 4.15

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has higher SiO2 contents and lower Ce/Pb ratios than the basanite(Table 1). If the alkali basalt could be produced by crustal contamina-tion of the basanite, its incompatible elements contents (especiallyLILE) would be higher than those of the basanite. This is opposite toour observations (Fig. 5). Therefore, no conspicuous crustal contami-nation occurred during the ascent and eruption of basaltic magmas.

5.2. Fractional crystallization

The Changle–Linqu basalts display a range of MgO contents asso-ciated with variations in Ni and Cr contents (Table 1). This, togetherwith the presence of olivine, pyroxene and plagioclase phenocrysts,suggests fractional crystallization during the magma ascent. Correla-tions between MgO/SiO2 and other elements in basaltic rocks may re-flect fractionation of some specific minerals. The alkali basalt exhibitsa negative correlation between SiO2 and MgO (Fig. 4d), suggestingolivine fractionation. However, SiO2 contents vary little whereasMgO contents show a large variation for the basanite, which is incon-sistent with the fractionation of olivine. The Ni contents are broadlynegatively correlated with (La/Sm)N ratios for the basanite (Table 1),suggesting the effect of partial melting because (La/Sm)N ratios arecontrolled by degree of melting and cannot be affected by fractionalcrystallization. Clinopyroxene may be one of the main fractionationphases because a positive correlation also exists between MgO andCaO/Al2O3 in the basalts (not shown). Significant plagioclase removaldid not happen because of the lack of negative Eu anomalies in thebasalts. In addition, the Sr–Nd–Pb isotope variations cannot be causedby fractional crystallization of anyminerals. Therefore, fractional crys-tallization alone is insufficient to account for the considerable varia-tions in both element and isotope compositions.

Basanite

07CL13 07CL03 07CL04 07CL05 07CL09

0.59 2.47 2.93 1.21 2.631.96 7.63 9.55 4.27 9.202.30 5.19 5.89 2.97 5.49

45.566 85.643 88.982 72.907 85.92717.461 18.262 18.221 18.094 18.2480.020 0.022 0.009 0.009 0.016

17.320 17.995 17.944 17.867 17.9810.033 0.061 0.064 0.052 0.062

15.418 15.503 15.454 15.451 15.4880.022 0.025 0.009 0.009 0.017

15.418 15.502 15.453 15.450 15.487150.923 264.401 290.120 256.544 300.74637.809 38.189 38.032 38.015 38.1550.024 0.025 0.010 0.008 0.018

37.660 37.927 37.744 37.761 37.857

Fig. 7. The O isotope composition of phenocryst minerals from Cenozoic basalts in theChangle–Linqu area, North China. (a) Phenocryst olivine, (b) phenocryst plagioclase,(c) O isotope fractionation between plagioclase and olivine. The δ18O range of mantleolivine is after Eiler et al. (1996), and the δ18O range of mantle plagioclase is calculatedfollowing the theoretical and empirical calibrations of Macpherson et al. (1998),Thirlwall et al. (1997), and Zheng (1993). The O isotope equilibrium fractionationlines between olivine and plagioclase are calculated after Zheng (1993).

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5.3. Partial melting of the mantle

Lithochemically, the basanite and alkali basalt could be generatedby partial melting of the same mantle source at relatively low andhigh degrees, respectively. However, there are significant differencesin initial Sr–Nd–Pb isotope compositions between the basanite andalkali basalt (Fig. 6), which is inconsistent with the same mantlesource. Therefore, we ascribe the two subtypes of basalts to origina-tion from different mantle sources.

During mantle melting, the trace element composition of basalticmelts is generally controlled by residual minerals (e.g., rutile andgarnet) and the bulk solid-melt partition coefficients (e.g., Yanget al., 2003). Thus, diagrams plotting element contents or ratiosversus partial melting index such as (La/Sm)n can be used to indicatethe presence of specific minerals as the residual phase. There is no ob-vious depletion of HFSE (Nb, Ta, Zr, and Hf) in the basalts, suggestingeither that Ti-bearing minerals such as rutile did not exist in the

residue or that they were very tiny in concentration if they did existbecause these accessory minerals are compatible for HFSE, especiallyNb and Ta (Prowatke and Klemme, 2005). This explanation is alsosupported by positive correlations between (La/Sm)n and Nb–Ta(Fig. 10a, b), suggesting that Nb and Ta were incompatible duringpartial melting. Garnet can cause large fractionation of HREE especial-ly in low-degree partial melting (e.g., Green et al., 2000). High Sm/Ybratios and low Lu/Hf ratios for the basanite (Table 1) may be attrib-uted to the presence of garnet as a residual phase during partialmelting.

5.4. Origin of low δ18O phenocrysts

Low δ18O values for phenocryst minerals are one of the importantcharacteristics for the Changle–Linqu basalts (Fig. 7). These may begenerated by three different mechanisms: high-T water–rock interac-tion, contamination by low δ18O crustal materials, and inheritancefrom the mantle sources (Zheng, 2012).

Water–rock interaction is ubiquitous when basaltic magmas areemplaced at shallow depths in the presence of water infiltratedfrom surface. Generally, low-T water–rock interaction would increasemineral δ18O values, whereas high-T water–rock interaction woulddecrease mineral δ18O values (Hoefs, 2009; Zheng, 2012). Meteoricgroundwater can infiltrate to depths of about 10 km and exchangeO isotopes with ambient rocks (Sheppard and Taylor, 1974). Thus, itis easy to make a genetic link between the low mineral δ18O valuesand the high-T water–rock interaction. In the δ18O–δ18O diagram,the alkali basalt displays a negatively correlated trend between oliv-ine and plagioclase (Fig. 7c), contrasting to the positively correlatedtrend for mineral O isotope equilibrium. Although the δ18O values ofplagioclase are only available from the two samples of basanite,they exhibit apparent O isotope temperatures of >1500 °C that aretoo high to represent the equilibrium temperature for continental ba-salts. In this regard, the high-T water–rock interaction is suggestedduring the basalt eruption. Because olivine is much more resistantto alteration by high-T fluid than plagioclase (e.g., Zheng and Fu,1998), the low δ18O values of olivine suggest its crystallization fromlow δ18O magmas (e.g., Zhang et al., 2009). Contamination by thecontinental crust with low δ18O values could be another possibilityfor the low δ18O olivines from the Changle–Linqu basalts, but thisprocess has been excluded as argued above. Therefore, the most pos-sible reason for the low δ18O olivines is that the mantle source ofbasaltic melts has relatively low δ18O values.

6. The nature of mantle source

6.1. Heterogenous juvenile SCLM sources

The Changle–Linqu Cenozoic basalts are characterized by the rela-tive enrichment of incompatible elements such as LILE and LREE(Fig. 5) but depletion of Pb and non-depletion of HFSE, with theOIB-like patterns of trace element distribution. This is different fromMORB-like patterns that exhibit the relative depletion of both LILEand LREE. The MORB is generally considered as originating frompartial melting of the normal asthenospheric mantle (e.g., Hofmann,1988; Salters and Stracke, 2004; Workman and Hart, 2005). In viewof the trace element composition, therefore, the normal asthenosphericmantle cannot be the source of OIB-like continental basalts (Zheng,2012). Lithospheric mantle, especially SCLM, has been considered oneof the main sources for continental basalts and other intraplate maficrocks (e.g., Barry et al., 2003; Graham et al., 2002; Hawkesworth et al.,1995; Liu et al., 2008; Wang et al., 2011; Zhang et al., 2009). In termsof both major and trace element compositions, the mantle sources forthe Changle–Linqu basalts appear to be fertile and enriched SCLM ratherthan refractory and depleted SCLM.

Fig. 8. Phenocryst olivine Ni, Ca, Mn contents and Fe/Mn ratios versus Fo values for Changle–Linqu basalts in North China, with reference to Herzberg (2011).

211Z. Xu et al. / Lithos 146-147 (2012) 202–217

The Changle–Linqu basalts exhibit relatively depleted Sr–Ndisotope compositions (Fig. 6a), with initial 87Sr/86Sr ratios of 0.7032to 0.7045 and εNd(t) values of −2.9 to 6.0 (Table 2). The alkali basalthas relatively low Na2O, 143Nd/144Nd and 206Pb/204Pb but high 87Sr/86Sr, whereas the basanite is opposite (Fig. 11), indicating heteroge-neous mantle sources for the two types of basalts. The youngest Ndmodel age is 334 Ma, which corresponds to a juvenile rather than

Fig. 9.Nb/U vs. Nb and Ce/Pb vs. Ce diagrams for Changle–Linqu basalts in North China. DataMORB and OIB are after Hofmann et al. (1986).

an ancient SCLM source. The juvenile SCLM can be produced by recenttransformation from the normal asthenospheric mantle, with geo-chemical inheritance in the Sr–Nd isotopes. With respect to the LILEand LREE, however, it requires crustal metasomatism for their enrich-ment (e.g., Wang et al., 2011; Zhang et al., 2009). In this regard, thejuvenile SCLM would be metasomatized by the crustally derivedmelts, generating the juvenile SCLM domains with compositional

for the upper and lower crusts are calculated after Rudnick and Gao (2003), and data for

Fig. 10. Ta and Nb vs. Th diagrams for Changle–Linqu basalts in North China.

212 Z. Xu et al. / Lithos 146-147 (2012) 202–217

heterogeneity. Themetasomatized SCLMdomainswould be fertile in themajor elements and thus susceptible to melting (Zheng, 2012).

6.2. Crustal component in mantle source

As argued above, the mantle sources of the Changle–Linqu basaltswould be generated by the interaction between the juvenile SCLMand the crustally derived melts. The crustal melts were derivedfrom partial melting of the crust that can be either the subductedcontinental/oceanic crust or the delaminated lower continental crust.Partial melting of the peridotites that are metasomatized by melts

Fig. 11. Plots of initial Sr–Nd–Pb isotope ratios versus Na2O contents for Changle–Linqubasalts in North China.

from the delaminated lower continental crust or carbonatitic liquidwas hypothesized to generate mafic magmas with the depletion ofHFSE (Liu et al., 2008; Zeng et al., 2010, 2011). However, this mecha-nism is inconsistent with the OIB-like patterns of trace element distri-bution in the Changle–Linqu basalts (Fig. 5). The mechanical mixturebetween the residual materials of melt-extracted crust and theasthenospheric mantle was suggested by Liu et al. (2008) to producethe OIB-like basalts. However, both the residual materials of melt-extracted crust and the MORB-like melt are characterized by therelative depletion of both LILE and LREE (Fig. 5), which is inconsistentwith the relative enrichment of LILE and LREE in the OIB-like basalts.

Low and variable δ18O values have been reported for UHP graniticgneiss and eclogite in the Dabie–Sulu orogenic belt, east-centralChina (Tang et al., 2008; Zheng et al., 2003, 2004, 2009). This orogenicbelt lies in the southeastern part of the NCC, so that there could be agenetic link between subduction of the continental crust and originof the Changle–Linqu basalts. Involvement of the subducted conti-nental crust in the mantle source could partly explain the featuresof major elements, Sr–Nd–Pb isotopes and O isotopes in these basalts.However, the subducted continental crust is characterized by the arc-like patterns of trace element distribution; so are post-orogenicmafic-ultramafic magmatic rocks in the Dabie–Sulu orogenic belt(Zhao and Zheng, 2009; Zhao et al., 2005, 2007a). This is not consis-tent with the Changle–Linqu basalts that exhibit the OIB-like patternsof trace element distribution.

Recycled oceanic crust in subduction zones has been consideredas one of the most important components in the mantle source ofOIB (e.g., Hofmann, 1997; Willbold and Stracke, 2006; Zindler andHart, 1986). Subducting oceanic crust would be metamorphicallydehydrated at shallow depths and lose considerable amounts ofwater-soluble elements such as Pb, Rb, Cs, Ba and LREE. In contrast,HFSE such as Nb and Ta are water-insoluble and thus would not betransported by aqueous fluids (Kogiso et al., 1997). Such fluidswould be produced by metamorphic dehydration in the rutile stabil-ity field and then react with the overlying mantle wedge peridotite,generating the mantle source for oceanic arc basalts with the relativeenrichment of LILE and LREE but the relative depletion of HFSE. Onthe other hand, the dehydrated oceanic crust would be partiallymelted at mantle depths of 100–300 km with rutile breakdown(Ringwood, 1990; Zheng, 2012), producing felsic melts with the rela-tive enrichment of LILE and LREE but the relative depletion of Pb andthe no depletion of HFSE. Such melts would react with the overlyingmantle wedge peridotite in oceanic subduction channels, generatingSi-deficient to Si-excess pyroxenites. Partial melting of the pyroxenitescan give rise to alkaline to tholeiitic basalts with the OIB-like traceelement distributions (Wang et al., 2011; Zhang et al., 2009; Zheng,2012). Oceanic gabbroic rocksmay suffer high-Twater–rock interactionduring their emplacement to acquire relatively low δ18O values (e.g.,Wang et al., 2003). The reaction of the juvenile SCLM peridotite withthe melt derived from the low δ18O oceanic crust can form the lowδ18O mantle source for low δ18O basalts. This is also supported by theoccurrence of low δ18O garnet pyroxenite xenoliths in Cenozoic basaltsfrom the Jiaohe area in the northeastern NCC, which are considered asremnants of the subducted oceanic crust (Yu et al., 2010).

Provided that themantle source for the Changle–Linqu basalts wasgenerated by the interaction between the juvenile SCLM peridotiteand the felsic melts derived from the subducting oceanic crust, it isimportant to determine where the oceanic crustal components camefrom. As indicated by seismic tomographic images, the Pacific platewas subducted below the eastern part of the NCC with the westwardthinned SCLM (Huang and Zhao, 2006; Xu and Zhao, 2009; Zhaoet al., 2007b). In this regard, the subduction of the Pacific oceaniccrust is the most possible candidate for the crustal components inthe mantle sources of Cenozoic continental basalts in eastern China.

The Changle–Linqu basalts have Pb isotope compositions (espe-cially 208Pb/204Pb ratios) similar to those of Indian MORB rather

Fig. 12. Fe/Zn vs. Mg# diagram for Changle–Linqu basalts in North China. Data forMORB are from PetDB (http://www.petdb.org/).

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than Pacific MORB (e.g., Martynov et al., 2012). Although there is noevidence for subduction of the Indian oceanic slab beneath the NCC,basaltic rocks with Indian MORB-like Pb isotope features have beenfound in Izu-Bonin and Kuril island arcs above western Pacific sub-duction zones (Straub et al., 2009). This suggests that the Pb isotopefeatures of western Pacific island arc basalts are dictated by recyclingof the crustal components in the oceanic subduction zone, with nogenetic link to subduction of the Indian oceanic crust. Therefore, itis reasonable to conclude that the Pacific oceanic crustal componentis incorporated into the mantle source of the Changle–Linqu basalts.

6.3. Lithology of the mantle source

Peridotite has been commonly considered as the mantle source ofoceanic and continental basalts. However, pyroxenite has also beensuggested for the mantle source of some basalts based on partialmelting experiments (Davis et al., 2011; Hirschmann et al., 2003;Keshav et al., 2004; Kogiso and Hirschmann, 2006; Kogiso et al.,2003) as well as whole-rock and olivine compositions (Herzberg,2011; Jackson and Dasgupta, 2008; Sobolev et al., 2005, 2007; Wanget al., 2011; Zhang et al., 2009). For example, Sobolev et al. (2005)suggested an olivine-free mantle source for Hawaii shield basaltsprimarily based on the Ni contents of olivine which are higher thanthose of olivine crystallized from melts derived from peridotitesource. Herzberg (2011) suggested a pyroxenite source for Hawaiiand Canary island basalts based on olivine and whole-rock major ele-ment compositions.

As shown in Table S2 and Fig. 8a, olivines with Fo between 80and 88 for the Changle–Linqu basanite and alkali basalt mostly haveFe/Mn ratios of >75, which are higher than those crystallized fromthe peridotite-derived melts as estimated by Herzberg (2011). It isknown that Fe/Mn partition coefficients (DFe/Mn) are lower than 1.0for clinopyroxene, orthopyroxene and garnet but greater than 1.0for olivine (Pertermann and Hirschmann, 2003; Walter, 1998). Thus,pyroxenite-derived melts would have higher Fe/Mn ratios thanthose of peridotite-derived melts. In this regard, the high olivine Fe/Mn ratios for the Changle–Linqu basalts suggest that their mantlesources are primarily composed of pyroxenite rather than peridotite.However, olivine Ni contents of these basalts are lower than thosecrystallized from melts derived from pure pyroxenite source (Fig. 8d),implying their derivation from the mantle sources with mixed litholo-gies between peridotite and pyroxenite (e.g., olivine pyroxenite andmetasomatized peridotite). Such non-peridotite sources were alsosuggested by Herzberg (2011) for Loihi OIB and some Canary OIB.

Basalt Fe/Zn ratios can be used as an indicator to infer the lithol-ogy of mantle sources. If a mantle source is pyroxenite rather thanperidotite, partial melts tend to have lower Fe/Zn ratios becauseclinopyroxene has higher DFe/Zn values than olivine (Le Roux et al.,2010, 2011). Most of the Changle–Linqu basalts have lower Fe/Znratios than MORB (Fig. 12), lending support to the conclusion thatthe pyroxene-rich lithologies would serve as the mantle source forthe Changle–Linqu basalts. The pyroxene-rich lithologies can be pro-duced by interaction between peridotite and felsic melts (e.g., Rappet al., 1999; Yaxley and Green, 1998). Thus, we speculate that theCenozoic basalts were derived from special mantle sources with thepyroxene-rich lithologies, which may be generated by interactionbetween the juvenile SCLM-wedge peridotite and the felsic meltsfrom the subducting Pacific oceanic crust in the Mesozoic. The crustalmetasomatism has been hypothesized to generate two types ofpyroxenites: Si-deficient pyroxenite produced by interaction with lessfelsic melts from metabasalt whereas Si-excess pyroxenite producedby interaction with more felsic melts from metasediment (Wang et al.,2011; Zhang et al., 2009).

Some experiments have indicated that melt–peridotite reactioncan consume olivine and consequently generate orthopyroxene(Rapp et al., 1999; Yaxley and Green, 1998). On the other hand,

Wang et al. (2010) found that melt–peridotite reaction can consumeolivine and generate clinopyroxene, attributing the difference to theCaO content of melt. In this regard, olivine that reacts with highand low CaO melts can generate clinopyroxene and orthopyroxene,respectively. If this inference is correct, the Si-deficient and Si-excess pyroxenites may have discrepancies not only in chemicalcharacteristics but also in mineral constituents. Mid-oceanic ridgebasalts and gabbros have higher CaO contents than oceanic sediment,with ~12 wt.% for mid-oceanic ridge basalts and gabbros but ~6 wt.%for global oceanic sediments (Hart et al., 1999; Klein, 2003; Plank andLangmuir, 1998). Thus the Si-deficient pyroxenite derived fromthe reaction between peridotite and melt of the oceanic metabasaltwould have more clinopyroxene than the Si-excess pyroxenitewhich was derived from the reaction between peridotite and melt ofthe oceanic metasediment. This feature may be used to explain somegeochemical differences between the basanite and alkali basalt inthis study. For example, the basanite has lower Fe/Zn ratios (Fig. 12),which may be attributed to more clinopyroxene in their mantlesource, because DFe/Zn for cpx is slightly higher than DFe/Zn for opx(Le Roux et al., 2010, 2011).

7. Implications for thinning of cratonic lithopsheric mantle inNorth China

Much progress has been made in understanding the diversity ofmagmatism above oceanic subduction zones, and the slab–mantle in-teraction is a key to generation of mantle sources in subduction chan-nels (e.g., Ringwood, 1990; Straub and Zellmer, 2012; Zheng, 2012).Major differences in basaltic magmatism between different subduc-tion zones are mainly caused by differences in thermally controlledpetrologic processes in subducting slabs. This is principally dictatedby variations in slab age and subduction dip (e.g., Gutscher et al.,2000; Hacker et al., 2003; Peacock and Wang, 1999; Rodríguez-González et al., 2012). In general, steep subduction tends to resultin significant release of aqueous fluids from subducting oceaniccrust, generating chloritized and serpentinized peridotites that serveas common mantle sources for arc magmatism at mantle depths ofabout 100 km (e.g., Ringwood, 1990; Syracuse and Abers, 2006;Tatsumi and Eggins, 1995). The oceanic arc basalts are thus derivedfrom partial melting of the fluid-altered peridotites, with the aqueousfluids originating from dehydration of the subducting oceanic crust inthe rutile stability field. The aqueous fluids are characterized by therelative enrichment of LILE, Pb and LREE but the depletion of HFSE,which is transferred to the altered peridotites and then to the oceanicarc basalts (Ringwood, 1990; Zheng, 2012).

In contrast, flat subduction tends to cause partial melting of theoceanic crust at mantle depths of 100–300 km (Beate et al., 2001;Gutscher et al., 2000; Ringwood, 1990). If both metabasalt andmetasediment partially melted, they gave rise to felsic melts that

Fig. 13. Tectonic sketches for the origin of Cenozoic continental basalts and its bearingon the SCLM thinning in the eastern part of North China. The slab–mantle interactionmodel of Ringwood (1990) has been advanced by Zheng (2012) to account for the gen-eration ofmantle sources for OIB-like continental basalts. (a) During the EarlyMesozoic,westward subduction of the Pacific plate would delaminate the lower part of ancientSCLM, with contemporaneous dehydration and melting of the oceanic crust at crustaland mantle depths, respectively. (b) During the Late Mesozoic, the juvenile SCLM hadreplaced the delaminated space for the ancient SCLM, with contemporaneous reactionof the overlying juvenile SCLM-wedge peridotite with adakitic to felsic melts from thesubducting Pacific oceanic crust. This generated different compositions of pyroxeniteat the lower part of the juvenile SCLM. (c) During the Cenozoic, partial melting of thepyroxenite together with peridotite took place to produce the continental basalts.

214 Z. Xu et al. / Lithos 146-147 (2012) 202–217

metasomatize the overlying mantle-wedge peridotite to form thepyroxenite-rich lithologies. Such ultramafic rocks would serve as thespecial mantle sources for OIB-like magmatism (Wang et al., 2011;Zhang et al., 2009). The continental basalts would thus originatefrom the melt-metasomatized peridotites in the SCLM, with themelt originating from partial melting of the subducting oceaniccrust with rutile breakdown. The melt is characterized by the relativeenrichment of LILE and LREE but the relative depletion of Pb andthe no depletion of HFSE, and this feature is transferred to themetasomatized peridotites and then to the OIB-like continentalbasalts (Ringwood, 1990; Zheng, 2012). Consequently, the mantlesource of fluid-altered origin is responsible for arc magmatismwhereas the mantle source of melt-metasomatized lithology is re-sponsible for OIB-like magmatism. It is the difference in the originof mantle sources that results in the difference in geochemical com-positions between their products of partial melting (Zheng, 2012).

With respect to the Cenozoic continental basalts in centralShandong, its geochemical features require the special mantlesources. In this regard, the subducting oceanic crust was metamor-phically dehydrated at first during eclogite-facies metamorphismand then partially melted at the mantle depths with the breakdownof rutile. It is known that the Pacific plate was subducted beneaththe eastern part of the NCC since the Mesozoic (e.g., Li and Li, 2007;Wu et al., 2005). The process of thinning the cratonic lithosphericmantle in North China may be associated with low-angle subductionof the Pacific plate beneath the eastern part of the NCC in the EarlyMesozoic (Fig. 13a). As a result, the lower part of ancient SCLM wasreplaced by the juvenile SCLM, with the consequent melt-peridotitereaction and thus the slab–mantle interaction above the oceanic sub-duction zone. The crustally derived melts are enriched not only inLILE and LREE, but also in radiogenic Sr–Nd isotopes with given vari-ations in O isotopes. On the other hand, they are depleted in Pb andno depletion of HFSE. Once they metasomatize the overlying juvenileSCLM-wedge peridotite, it generates both Si-deficient and Si-excesspyroxenites in the Late Mesozoic (Fig. 13b). The breakdown of rutileis a key to the no depletion of HFSE in the refertilized SCLM domains.The pyroxene-rich lithologies were storied at the lower part of SCLMfor about one hundred million years, allowing preservation of thelow δ18O signature in the mantle sources. In the Cenozoic, it becamepartially melted with garnet as a residual phase, producing the conti-nental basalts as observed today in the eastern part of the NCC(Fig. 13c). In this context, the Cenozoic basalts are a snapshot of theslab–mantle interaction during the SCLM thinning in North China.

In order to reconcile the thermal state of subducting slabs withgeochemical observations, Gutscher et al. (2000) suggested a three-stage model for thermal evolution of oceanic subduction zones:(1) steep subduction produces a narrow calc-alkaline arc, typically~300 km from the trench, above the asthenospheric mantle wedge;(2) once flat subduction begins, the lower plate travels several hundredkilometers at nearly the same depth, thus remaining in a P–T windowallowing slab melting over this broad distance; and (3) once flatsubduction continues for several million years, the asthenosphericmantle wedge disappears, and a volcanic gap results. The geophysicalobservations show a gradual reduction of the SCLM thickness fromwest to east in the eastern part of the NCC (Chen, 2009; Huang andZhao, 2009; Tian et al., 2009; Xu et al., 2009). This is ascribed to westernsubduction of the Pacific plate in the Mesozoic (Zhang et al., 2009; Zhuand Zheng, 2009). In this regard, subduction of the Pacific plate hasplayed the first-order role in thinning the cratonic lithopsheric mantleinNorth China (ZhengandWu, 2009). In associationwith this tectonism,there would be not only the mechanical delamination of the lower partof SCLM but also the thermo-chemical erosion along the slab interfaceby the upwelling asthenospheric mantle. The melt–peridotite reactionis responsible for the Mesozoic interaction between the juvenile SCLMand the subducting Pacific oceanic crust, whereas its final derivative isthe Cenozoic continental basalts in the easternmargin of NCC. Therefore,

the westward subduction of the Pacific plate has caused both physicaland chemical changes of the SCLM in the NCC.

8. Conclusions

The Cenozoic continental basalts in central Shandong of NorthChina exhibit the OIB-like patterns of trace element distribution andthe depleted Sr–Nd isotope compositions. This indicates that theirmantle sources are a kind of the juvenile SCLM with considerableheterogeneities in geochemistry. Themantle sourceswould have formed

215Z. Xu et al. / Lithos 146-147 (2012) 202–217

by reaction of the juvenile SCLM-wedge peridotite with felsic meltsderived from the subducting Pacific oceanic crust in the Mesozoic. Bythis way, westward subduction of the Pacific plate beneath the easternpart of the NCC would have served as a geodynamic mechanism to de-laminate the lower part of the ancient SCLM, resulting in destruction ofthe North China Craton. Mineral phenocrysts from the basalts primarilyexhibit lower δ18O values than the normalmantle value. This is not a re-sult of high-T water–rock interaction or low δ18O crustal contaminationduringmagma emplacement. Instead, it is attributed to O isotope inher-itance from their mantle sources that would become depleted in 18O bythe melt-peridotite reaction via the slab–mantle interaction duringthe oceanic subduction in the Mesozoic. Olivine element compositionssuggest pyroxene-rich lithologies rather than olivine-rich lithologiesfor the continental basalts. The different compositions of pyroxenitesand metasomatized peridotites are hypothesized to form through thereaction of the juvenile SCLM-wedge peridotite with the felsic meltsderived from partial melting of the subducting Pacific oceanic crustwith the rutile breakdown. Partial melting of the fertile and juvenileSCLM domains in the Cenozoic is proposed to give rise to the continentalbasalts. Therefore, the Cenozoic continental basalts provide a snapshot ofthe slab–mantle interaction in association with the SCLM thinning inNorth China due to the Mesozoic flat subduction of the Pacific plate.

Acknowledgments

This study was supported by funds from the Chinese Academyof Sciences (KZCX2-YW-Q08) and the Natural Science Foundationof China (91014007 and 40921002). Thanks are due to J. Tang andQ.-L. Yang for their assistance with the field sampling, to X.-P. Zhaand B.-C. Ding for their assistance with O isotope analyses, to H.Qian, P. Xiao, J.-F. He, H.-Y. Zhou, G.-Z. Li and H. Guo for their assis-tance with Sr–Nd–Pb isotope analyses, and to S. Zheng for assistancewith olivine composition analyses. We are grateful to the editorN. Eby and two anonymous reviewers for their comments that helpedimprove the presentation.

Appendix A. Supplementary data

Supplementary data to this article can be found online at http://dx.doi.org/10.1016/j.lithos.2012.05.019.

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