Origin of low δ26Mg Cenozoic basalts from South China Block and their geodynamic implications

Accepted Manuscript Origin of low δ 26Mg Cenozoic basalts from South China Block and their geodynamic implications Jian Huang, Shu-Guang Li, Yilin Xiao, Shan Ke, Wang-Ye Li, Ye Tian PII: DOI: Reference:

S0016-7037(15)00268-9 http://dx.doi.org/10.1016/j.gca.2015.04.054 GCA 9260

To appear in:

Geochimica et Cosmochimica Acta

Received Date: Accepted Date:

13 November 2014 30 April 2015

Please cite this article as: Huang, J., Li, S-G., Xiao, Y., Ke, S., Li, W-Y., Tian, Y., Origin of low δ 26Mg Cenozoic basalts from South China Block and their geodynamic implications, Geochimica et Cosmochimica Acta (2015), doi: http://dx.doi.org/10.1016/j.gca.2015.04.054

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Origin of low δ26Mg Cenozoic basalts from South China Block and their geodynamic implications Jian Huanga, Shu-Guang Lib,a*, Yilin Xiaoa*, Shan Keb, Wang-Ye Lia, Ye Tiana a

CAS Key Laboratory of Crust-Mantle Materials and Environments, School of Earth and

Space Sciences, University of Science and Technology of China, Hefei 230026, China b

State Key Laboratory of Geological Processes and Mineral Resources, China University of

Geosciences, Beijing 100083, China

*Corresponding authors: E-mail: [email protected]; [email protected]

Revised version submitted to GCA (2015.04.19) Abstract: 377 words Main text: ~6460 words Figures: 13 Tables: 2 Supplementary Table: 1

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Abstract Origin of low δ26Mg basalts is a controversial subject and has been attributed to interaction of isotopically light carbonatitic melts derived from a subducted oceanic slab with the mantle (Yang et al., 2012), or alternatively, to accumulation of isotopically light ilmenite (FeTiO3) in their mantle source (Sedaghatpour et al. 2013). To study the origin of low δ26Mg basalts and evaluate whether Mg isotope ratios of basalts can be used to trace deeply recycled carbon, high-precision major and trace element and Mg isotopic analyses on the Cenozoic alkaline and tholeiitic basalts from the South China Block (SCB), eastern China have been carried out in this study. The basalts show light Mg isotopic compositions, with δ26Mg ranging from –0.60 to –0.35‰. The relatively low TiO2 contents (<2.7 wt.%) of our basalts, roughly positive correlations between δ26Mg and Ti/Ti* and their constant Nb/Ta ratios (16.4–20) irrespective of variable TiO2 contents indicate no significant amounts of isotopically light ilmenite accumulation in their mantle source. Notably, the basalts display negative correlations between δ26Mg and the amounts of total alkalis (i.e., Na2O + K2O) and incompatible trace elements (e.g., Ti, La, Nd, Nb, Th) and trace element abundance ratios (e.g., Sm/Yb, Nb/Y). Generally, with decrease of δ 26Mg values, their Hf/Hf* and Ti/Ti* ratios decrease, whereas Ca/Al and Zr/Hf ratios increase. These features are consistent with incongruent partial melting of an isotopically light carbonated mantle, suggesting that large variations in Mg isotope ratios occurred during partial melting of such carbonated mantle under high temperatures. The isotopically light carbonated mantle were probably formed by interaction of the mantle with low δ26Mg carbonatitic melts derived from the deeply subducted low δ26Mg carbonated eclogite transformed from carbonate-bearing oceanic crust during plate subduction. As only the Pacific slab has an influence on both the North China Block (NCB) and SCB, our results together with the study of Yang et al. (2012) demonstrate that the recycled carbonatitic melts might have originated from the stagnant Pacific slab beneath East Asia in the Cretaceous and Cenozoic and that a widespread carbonated upper mantle exists beneath eastern China, which may serve as the main source for the <110Ma basalts in this area. Thus, our study demonstrates that Mg isotope ratios of basalts are a powerful tool to trace deeply recycled carbon.

Key words: Magnesium isotopes; Deep carbon cycle; Basalts; Pacific slab subduction; Carbonated mantle.

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1.

Introduction Magnesium (Mg) is a major constituent of the Bulk Silicate Earth (2.8 wt.%, Rudnick

and Gao, 2003) and the second most important cation in the seawater (~0.13 wt.%, Millero, 1974). It has three stable isotopes, 24Mg, 25Mg and 26Mg, with natural abundances of 78.99%, 10.00% and 11.01%, respectively (Rosman and Taylor, 1998). The relatively mass difference between 24Mg and 26Mg is ~8%, large enough to produce significant Mg isotope fractionation in cosmochemical, geochemical and biological processes. Mg isotope fractionation has been used to play constraints on a variety of scientific issues such as the evolution of the early solar system (e.g., Lee et al., 1976; Bizzarro et al., 2004), equilibrium temperatures of metamorphic and mantle rocks (e.g., Li et al., 2011; Huang et al., 2013), continental weathering (e.g., Tipper et al., 2006; Teng et al., 2010a; Huang et al., 2012a; Liu et al., 2014; Wimpenny et al., 2014a), plant growth (e.g., Black et al., 2006; Bolou-Bi et al., 2010), enzyme synthesis (Buchachenko et al., 2008) and paleoclimate changes (Saenger and Wang, 2014). With respect to the cosmochemistry and geochemistry of Mg isotopes, one of the most important findings is that some high-Ti lunar basalts and some terrestrial basalts have very light Mg isotopic compositions with δ26Mg as low as –0.60 (Yang et al., 2012; Sedaghatpour et al., 2013). The interpretations for the origins of the low δ26Mg basalts are inconsistent. Because carbonate rocks (e.g., dolostone and limestone) and minerals (e.g., calcite, aragonite, dolomite and magnesite) have extremely light Mg isotopic compositions, with δ 26Mg varying widely from –5.54 to –0.47 (e.g., Galy et al., 2002; Young and Galy, 2004; Tipper et al., 2006; Pogge von Strandmann et al., 2008a; Higgins and Schrag, 2010; Jacobson et al., 2010; Ke et al., 2011; Pokrovsky et al., 2011; Wombacher et al., 2011), the low δ26Mg terrestrial basalts from the North China Block (NCB), eastern China have been suggested to result from interaction of their mantle source with isotopically light carbonatitic melts derived from the subducted oceanic slab (Yang et al., 2012). However, based on the large Mg isotope fractionation recorded in high-Ti and low-Ti lunar basalts, with the former generally having much lower δ26Mg values than the latter (–0.59 to –0.37 vs. –0.33 to –0.02, Sedaghatpour et al., 2013), the low δ26Mg high-Ti lunar basalts have been suggested to originate from an isotopically light mantle source produced by crystallization of ilmenite (FeTiO3) with low δ26Mg at the late stage in the lunar magma ocean (Sedaghatpour et al., 2013). To evaluate wheather Mg isotope ratios of basalts can be used to trace deeply recycled carbon, it is necessary to re-evaluate whether the low δ26Mg basalts from eastern China were caused by mixture of isotopically light carbonatitic melts into their mantle source as suggested by Yang 3

et al. (2012), or alternatively, by accumulation of isotopically light ilmenite in their mantle source (Sedaghatpour et al., 2013), although it is not sure whether ilmenite has a light Mg isotopic composition, because so far no Mg isotopic data for it has been reported,. Here, we present high-precision Mg isotopic analyses on a suite of well-characterized Cenozoic alkaline and tholeiitic basalts from the South China Block (SCB), eastern China. The basalts show a large variation in major and trace element geochemistry as well as similar depleted Sr-Nd isotopic compositions (e.g., Zou et al., 2000; Chen et al., 2009; Wang et al., 2011). Previous studies have demonstrated that a carbonated mantle may be the main source for the alkaline basalts (Chen et al., 2009; Zeng et al., 2010; Yang et al., 2012; Sakuyama et al., 2013). Our results show that these basalts have low δ26Mg values. Interestingly, a negative correlation between δ26Mg and TiO2 exists in the basalts studied here. Thus, the Mg isotopic compositions of these basalts firstly allow us to re-evaluate whether the low δ26Mg basalts from eastern China were caused by recycled carbonates through oceanic plate subduction or accumulation of isotopically light ilmenite in their mantle source. Secondly, we refer the geodynamic implications deduced from the low δ26Mg basalts from eastern China.

2.

Geological settings and sample descriptions

In eastern China, the Cenozoic volcanic rocks are widely distributed along the coastal provinces and adjacent offshore shelf extending over 4000 km from Heilongjiang Province in the north to Hainan island in the south in the eastern edge of the Eurasian continent (Fig. 1). They constitute an important part of the volcanic belt of the western circum-Pacific rim and are one of the world’s presently active tectono-magmatic regions (e.g., Zhou and Armstrong, 1982). The Cenozoic volcanic rocks are mainly alkaline basalts that are thought to represent melts derived from the upper mantle, given their depleted Sr-Nd isotopic compositions and OIB-like trace-element signatures in spidergram (e.g., enrichment in Nb, Ta and LREEs, and negative K and Pb anomalies; e.g., Zhou and Armstrong, 1982; Peng et al., 1986; Liu et al., 1994; Zou et al., 2000; Xu et al., 2005; Tang et al., 2006; Chen et al., 2009; Zeng et al., 2010, 2011; Wang et al., 2011). The samples investigated in this study were collected from Pingmingshan, Anfengshan, Fangshan, Chongren, and Longyou from the SCB (Fig. 1). K-Ar dating results show that basalts from Pingmingshan and Anfengshan have ages of 7.3‒12.3 and 4.0‒6.4 Ma, respectively (Chen and Peng, 1988; Jin et al., 2003), whereas those from Fangshan, Chongren and Longyou have ages of 2.9‒3.5, ~26.4, and 9.0‒9.4 Ma, respectively (Chen and Zhi, 1988; 4

Ho et al., 2003). Mantle peridotite xenoliths are common in all localities (e.g., Qi et al., 1995; Jin et al., 2003; Reisberg et al., 2005). Twenty-three basaltic samples were selected for investigations, and all of them are of the porphyritic texture. Most of the studied samples are fresh and unaltered, exceptions are samples 13AFS9-10 from Anfengshan, 10LYSK11 from Longyou, which are altered with iddingsitization of olivine phenocrysts. The phenocrysts consist predominantly of olivine in the Anfengshan basalts, of olivine and clinopyroxene in the Fangshan and Longyou basalts, and of olivine, clinopyroxene and plagioclase in the Pingmingshan and Chongren basalts. The groundmass in these basalts is variable and mainly consists of plagioclase, olivine, augite, nepheline, magnetite and glass.

3. Analytical methods

3.1. Major and trace element analysis The samples were sawed into slices and only central fresh parts were used for bulk-rock analyses. The pieces were crushed in a corundum jaw crusher to 60 mesh, and then ~60 g of each crushed sample was powdered in an agate ring mill to <200 mesh in size. Bulk rock abundances of major elements were determined using an X-ray fluorescence spectrometer (XRF) on glass disks at the laboratory of ALS minerals at Guangzhou. Pre-ignition was used to determine the loss on ignition (LOI) prior to major elements analyses. Accuracy and precision for major oxides are generally better than 1% based on replicate analyses of certified USGS rock standards. Bulk rock trace element data were obtained by an ELAN DRCII inductively coupled plasma-mass spectrometry (ICP-MS) at the University of Science and Technology of China (USTC) after ultrapure acid digestion (HNO 3 + HF + HClO4) of sample powders (~50mg) in Teflon bombs. Analytical procedures were described in detail by Huang et al. (2012b, 2014). The measured values of international USGS standards (BHVO-2 and BIR-1) are in satisfactory agreement with the recommended values within error, and the precision and accuracy for majority of trace elements analyzed are better than 6% (Table S1 in Supplementary Materials).

3.2. Mg isotope analysis

Magnesium isotopic analyses were performed at the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences (Beijing) (CUGB), 5

following the procedures very similar to those established by Yang et al. (2009), Teng et al. (2010b) and Teng and Yang (2014). A brief description is given below. All chemical procedures were conducted in a clean laboratory environment at CUGB. Sample powders were dissolved in Savillex screw-up beakers in a mixture of concentrated HF–HCl–HNO3. Chemical separation of Mg was achieved by cation exchange chromatography with Bio-Rad 200–400 mesh AG50W-X8 pre-cleaned resin in 1 N HNO3 media. The same column procedure was processed twice for all samples in order to obtain a pure Mg solution for mass spectrometry and to check the efficiency of our column to separate Mg from interference cations. The eluted solutions were firstly evaporated to dryness and then re-dissolved in 3% HNO3, ready for final dilution immediately prior to analysis. Three USGS reference materials (e.g., BHVO-2, AGV-2 and GSP-2) were processed through column chemistry with each batch of the investigated samples. The total procedural blanks during the course of this study were ~8 ng, comparable to that of Teng et al. (2010b). Magnesium isotopic compositions were measured by the sample-standard bracketing method on a Neptune Plasma MC-ICPMS in a low-resolution mode. The in-run precision on the 26Mg/24Mg ratio for a single block of 40 ratios is < ± 0.02‰ (2SD). The internal precision on the measured

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Mg/24Mg ratio based on 4 repeated analyses of the same sample solution

during analytical sessions of this study, is ≤ ±0.08‰ (2SD, Table. 2). The results are reported in the conventional δ notation that is defined as δ XMg = [(XMg/24Mg)sample/(XMg/24Mg)DSM3 – 1] × 1000, where X = 25 or 26, and DSM3 is an international reference of Mg solution made from pure Mg metal (Galy et al., 2003). Analyses of the well-characterized USGS Mg rock standards in this study yielded δ26Mg = –0.25 ± 0.05‰ (2SD, n=2) for BHVO-2, –0.15 ± 0.03‰ (2SD, n=2) for AGV-2 and –0.03 ± 0.09‰ for GSP-2 (Table 2), which are in excellent agreement with previously published values within error (e.g., Teng et al., 2007, 2010a,b; Huang et al., 2009; An et al., 2014; Teng and Yang, 2015).

4. Results Results for bulk-rock major and trace element concentrations and Mg isotopic compositions of the studied basalts are listed in Tables 1 and 2, respectively.

4.1. Major and trace elements

The investigated samples have a large compositional range of SiO2 (40.6 to 51.1 wt.%) and high contents of MgO (6.91–13.1 wt.%) and TiO2 (1.81–2.64 wt.%). They also show high 6

total alkalis (Na2O + K2O = 3.71–9.06 wt.%), with Na2O/K2O ratios ranging from 1.4 to 3.3, indicating their alkali-rich and high-sodium nature. Following the nomenclature of La Bas et al. (1986), samples from Anfengshan, Pingmingshan and Longyou are basanites, those from Fangshan are trachybasalt, and those from Chongren are normal basalts (Fig. 2). All basalts are alkaline except for samples from Chongren, which are tholeiites (Fig. 2). No clear correlation was observed between MgO and SiO2, TiO2, Fe2O3t, CaO and K2O, while Al2O3 displays a slightly negative correlation with MgO (Fig. 3). In the chondrite-normalized REE diagram (Fig. 4a), all basalts show enrichment of LREEs over HREEs ([La/Yb]N = 6.0 to 32), with no significant Eu or Ce anomaly. In primitive mantle-normalized spidergram (Fig. 4b), all but rocks from Chongren resemble many ocean island basalts in terms of enrichment in Nb and Ta, and depletion in K and Pb relative to LREEs. Additionally, the basalts generally show negative Zr, Hf and Ti (Hf/Hf* = 0.49–0.86, Ti/Ti* = 0.32–1.03, indexes for Hf and Ti anomalies, respectively) and high Ca/Al ratios (0.47–0.81) (Fig. 5). Nb/U and Ce/Pb ratios of the Cenozoic basalts range from 36 to 52 and 13 to 34, respectively (Table 1), similar to those of MORBs and OIBs (Nb/U = 47 ± 10, Ce/Pb = 25 ± 5, Hofmann et al, 1986), but much higher than those of continental crust (Nb/U = 6.2, Ce/Pb = 4, Rudnick and Gao, 2003).

4.2. Magnesium isotopes In the plot of δ25Mg vs. δ26Mg (Fig. 6a), the basalts and the USGS standards fall along the terrestrial equilibrium mass fractionation line with a slope of 0.521, similar to previous studies of natural and synthetic samples (e.g., Young and Galy, 2004; Yang et al., 2009, 2012; Li et al., 2010; Teng et al., 2010a,b). The Δ25Mg´ (Δ25Mg´ = δ25Mg´ – 0.521δ26Mg´, where δXMg´ = 1000 × ln[(δXMg + 1000)/1000], X = 25 or 26, Young and Galy, 2004) values for all basalts range from –0.01 to 0.04, very close to zero (Table 2). In general, our Cenozoic basalts have highly variable but overall light Mg isotopic compositions, with δ 26Mg values ranging from –0.60 to –0.35‰ (Fig. 6). These δ26Mg values are similar to those of <110 Ma basalts from the NCB (–0.60 to –0.42‰, Yang et al., 2012) but obviously lighter than that of the normal mantle (~ –0.25, Teng et al., 2007, 2010b; Handler et al., 2009; Bourdon et al., 2010) and >120Ma basaltic rocks from the NCB (–0.27 ± 0.05‰, n=5, 2SD, Yang et al., 2012) (Fig. 6b). Interestingly, the Mg isotopic compositions of the basalts are negatively correlated with the abundance of total alkalis (Na2O + K2O) and incompatible elements (e.g., La, Nd, Ti, Nb, Th) as well as trace element abundance ratios of Sm/Yb and Nb/Y (Figs. 7 and 8).

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5. Discussion

5.1. Effects of shallow-level processes 5.1.1. Post-magmatic alteration Loss on ignition (LOI) of our basalts highly vary from –0.08 to 4.96 wt.%, suggesting that some samples have experienced alterations, manifested by the transformation of olivine phenocrysts to low-T iddingsite in samples 13AFS9-10 and 10LYSK11. However, good correlations between abundances of fluid-mobile elements (e.g., Ba, Sr, Pb, Th, U, La, Nd) and the fluid-immobile element Nb (Fig. 9) suggest that the effect of alteration, which would markedly disturb such correlations, is not significant. In addition, residual products of basalts (e.g., saprolite and soil) usually have heavier Mg isotopic compositions compared to their protoliths due to preferential incorporation of

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Mg into the secondary Mg-bearing clay

minerals (e.g., Pogge von Strandmann et al., 2008b; Teng et al., 2010a; Huang et al., 2012a; Liu et al., 2014; Wimpenny et al., 2014b). However, the studied basalts have lighter Mg isotopic compositions relative to normal mantle-derived magmatic rocks (e.g., MORBs and OIBs, –0.25 ± 0.07, Teng et al., 2010b) (Figs. 6-8), opposite to the expected results of alteration. Therefore, the low δ26Mg values of our basalts are unlikely to result from post-magmatic alteration.

5.1.2. Crustal contamination The Cenozoic alkaline basalts from eastern China have been suggested to originate from the upper mantle with negligible crustal contamination based on previous geochemical and Sr-Nd isotopic studies (e.g., Zhou and Armstrong, 1982; Peng et al., 1986; Liu et al., 1994; Zou et al., 2000; Xu et al., 2005; Tang et al., 2006; Chen et al., 2009; Zeng et al., 2010, 2011; Wang et al., 2011). This interpretation can also be applied to the alkaline basalts investigated here because their OIB-like features (e.g., positive Nb and Ta anomalies and negative K and Pb anomalies, Fig. 4b) and depleted Sr-Nd isotopic compositions (Zou et al., 2000; Chen et al., 2009; Wang et al., 2011) are inconsistent with crustal contamination. The occurrence of mantle xenoliths indicates that the basaltic magmas ascended rapidly, which would leave little time for magma evolution or wall-rock assimilations. OIB-like Ce/Pb and Nb/U ratios of our basalts (Table 1) further indicate negligible crustal contamination that would lower these ratios of the basaltic magmas (Hofmann et al., 1986; Sun and McDonough, 1989). Finally, the negative correlation between SiO2 and Na2O + K2O contents (Fig. 2) is also inconsistent with crustal contamination that would result in an opposite trend. 8

5.1.3. Crystal fractionation Most of the investigated basalts have relatively high Mg# (0.64 to 0.74, molar Mg/[Mg + Fe+2], Table 1), suggesting that their compositions are close to those of the primary magmas (Langmuir et al., 1977) and that insignificant fractional crystallization of olivine and pyroxene occurred. A few basalts (i.e., 13AFS1-4 and 10CR1-2) with relatively low Mg# (0.55 to 0.60) and Ni (<170 ppm) and Cr (<170 ppm) contents (Table 1) might have experienced olivine and pyroxene fractionation. However, previous studies show that fractional crystallization of olivine and pyroxene cause no detectable Mg isotope fractionation during basalt differentiation (e.g., Teng et al., 2007), but can result in increases in incompatible element contents of basalts (Allègre et al., 1977). Thus, if fractional crystallization of these two minerals is significant, no correlations between δ26Mg values and incompatible element contents should be observed in basalts. Our basalts display obviously negative correlations between δ26Mg values and contents of incompatible elements, such as Ti, La, Nd, Nb and Th (Fig. 7), suggesting negligible fractional crystallization of olivine and pyroxene. Furthermore, since fractional crystallization of minerals such as olivine and pyroxene has a very limited effect on two incompatible element abundance ratios (e.g., Sun and Davis, 1975; Minster and Allègre, 1978), the negative correlations between δ26Mg values and Sm/Yb and Nb/Y ratios in our basalts (Fig. 8) must record the original features of the primary magmas. Finally, no negative Eu anomalies (Fig. 4a) are indicative of negligible removal of plagioclase. Thus, the observed variations in geochemistry and δ 26Mg values of our basalts are largely related to mantle processes.

5.2. Geochemical variations with degree of partial melting in the upper mantle

Since shallow-level processes (including post-magmatic alteration, crustal contamination and crystal fractionation) have insignificant effects, the compositions of the studied basalts are close to the primary magmas. Thus, the different features in geochemistry are likely to result from different mantle sources or from different degrees of partial melting of similar mantle sources. Given that the Sr-Nd isotopic compositions of these basalts are depleted and vary in a narrow range (Zou et al., 2000; Chen et al., 2009; Wang et al., 2011), their LILE and LREE enrichments suggest that they were probably derived from similar sources, which were recently enriched by mantle metasomatism after a long-term depletion. Since La and Sm are more incompatible than Yb during partial melting of the mantle rocks, La/Yb and Sm/Yb ratios are sensitive to partial melting degree. As illustrated in Fig. 10, variable degrees 9

(2%–20‰) of batch melting of a hypothetical light REE-enriched mantle source ([La/Yb]n >1) in the garnet stability field can produce the La/Yb vs. Sm/Yb systematics of the alkaline and tholeiitic basalts investigated here. Specifically, alkaline basalts with high La/Yb and Sm/Yb from Pingmingshan ratios have the lowest degree of partial melting (~2%), while tholeiitic basalts with lower La/Yb and Sm/Yb ratios from Chongren have the highest degree of melting (~20%). These estimates are consistent with the different contents of incompatible elements such as Ba, Th, U, La, Sr, Nd and Nb (Fig. 9), because the amount of incompatible elements in basaltic melts increases with decreasing degree of partial melting. In addition, the Cenozoic basalts investigated here have highly variable Nb/Y ratios that are also strongly dependent on the degree of partial melting, because Nb is much more incompatible than Y during partial melting of the mantle rocks (Fig. 8b). Therefore, the data and discussions above lead us to suggest that the basalts studied here originated from the upper mantle and their geochemical differences closely reflect variable degrees of partial melting. With decreasing degree of partial melting, the basalts display higher total alkalis and TiO2 contents, higher concentrations of incompatible elements (e.g., Ba, Sr, La, Nd, Nb, U, Th), increased La/Yb, Sm/Yb and Nb/Y ratios, more pronounced negative K, Zr, Hf and Ti anomalies in primitive mantle (PM)-normalized trace element spidergram, and much lighter Mg isotopic compositions, as shown in Figs. 4 and 7-10. 5.3. Origin of the low δ26Mg Cenozoic basalts from the SCB

Previous studies have shown that no significant Mg isotope fractionation occurs during partial melting of the mantle and subsequent basalt differentiation as well as granite differentiation (Teng et al., 2007, 2010b; Handler et al., 2009; Bourdon et al., 2010; Li et al., 2010; Liu et al., 2010; Telus et al., 2012), implying that basalts and granites theoretically should have a mantle-like Mg isotopic composition if no isotopically distinct components (such as sedimentary rocks, including carbonate rocks, shale, loess and soil, which have δ26Mg ranging from –5.57 to 1.8, e.g., Galy et al., 2002; Young and Galy, 2004; Tipper et al., 2006; Pogge von Strandmann et al., 2008a; Higgins and Schrag, 2010; Jacobson et al., 2010; Li et al., 2010; Teng et al., 2010a; Wombacher et al., 2011; Liu et al., 2014) are involved during their genesis. The heavy Mg isotopic compositions (δ26Mg = up to 0.44) of I-type granites from southern California was attributed to the recycled high δ26Mg sedimentary rocks in their source (Shen et al., 2009). Yang et al. (2012) has identified a suite of low δ26Mg continental basalts from the NCB and interpreted such feature as due to interactions between 10

the mantle and the isotopically light carbonatitic melts derived from the subducted oceanic slab. On the other hand, Sedaghatpour et al. (2013) more recently reported that high-Ti lunar basalts also display light Mg isotopic compositions (as low as –0.60). Based on the negative correlation between δ26Mg and TiO2 in high-Ti basalts, they suggested that ilmenite has light Mg isotopic compositions and the accumulation of ilmenite in the mantle source at the late stage in the lunar magma ocean shifts high-Ti basalts to low δ26Mg values (Sedaghatpour et al., 2013). Similarly, a negative correlation between δ26Mg and TiO2 has also been observed in our basalts (Fig. 7b). Therefore, we have to evaluate the possibility that the light Mg isotopic compositions of our basalts might have resulted from the ilmenite accumulation in their mantle source. Several lines of evidence are against this possibility. First, our basalts have TiO2 contents (<2.7 wt.%) much lower than those of high-Ti lunar basalts (>6 wt.%, Sedaghatpour et al., 2013), implying that no significant accumulation of ilmenite in the mantle source for our basalts. Second, as ilmenite generally displays an enrichment of Nb-Ta and preferentially incorporates Ta relative to Nb with DTa/DNb = ~1.3 between ilmenite and mafic melts (Dygert et al., 2013), accumulation of ilmenite in the mantle source would cause a negative correlation between Nb/Ta and TiO2 rather other constant Nb/Ta ratios in our basalts irrespective of variable TiO2 (Fig. 11). Third, roughly positive correlation between δ26Mg and Ti/Ti* (Fig. 13d) in our basalts also doesn’t stand for this interpretation, because if large amounts of isotopically light and Ti-rich ilmenites are present in the mantle source, the opposite should be observed. Thus, the light Mg isotopic compositions of the investigated basalts cannot be resulted from ilmenite accumulation in their mantle source. In addition to the negative correlation between δ26Mg and TiO2, the δ26Mg values of our basalts also decrease with the amounts of other incompatible elements (e.g., La, Nd, Nb, Th) and trace element abundance ratios (e.g., Sm/Yb, Nb/Y) (Figs. 7, 8) that are sensitive to partial melting. This suggests that large variations in Mg isotope ratios have occurred during partial melting of the mantle under high temperatures and pressures, with melts produced by low degrees of melting having lighter Mg isotopic compositions relative to melts produced by high degrees of melting. A dissolution experiment on Boulder Creek Granodiorite has shown that concomitant variations in δ26Mg values of reactive fluids reflect conservative mixing of Mg released from isotopically distinct minerals (e.g., chlorite, biotite and hornblende) rather than Mg isotope fractionation. It is experimentally determined that during partial melting of carbonated peridotite, the first incipient melts near solidus are carbonatitic melts, which evolve to carbonated silicate melts (≤ 25 wt. % CO2 in melts) with increasing degree of 11

melting (Dasgupta et al., 2007, 2013). Hence, it is expected that during this evolution path, there should exist large variations in Mg isotopic compositions because of different ratios of isotopically light carbonatitic melts/isotopically heavy silicate melts at different degrees of partial melting of carbonated peridotite. Teng et al. (2007, 2010b) demonstrated that Mg isotope fractionation is insignificant during partial melting of dry peridotite and subsequent basalt differentiation. In the studies of Teng et al. (2007, 2010b), the Mg-rich minerals involved during magmatic processes are olivine and pyroxene which usually have mantle-like Mg isotopic compositions (e,g., Handler et al., 2009; Yang et al., 2009; Huang et al., 2011; Liu et al., 2011; Pogge von Strandmann et al., 2011; Xiao et al., 2013). This may explain why no significant Mg isotope fractionation was observed in their studies. The basalts in the present study overall have much lighter Mg isotopic compositions compared to global mid-ocean ridges basalts (MORBs, Figs. 7, 8 and 13), suggesting that the mantle source for the studied basalts is not a dry garnet peridotite that has an average Mg isotopic composition identical to that in MORBs (δ26Mg = –0.26 ± 0.07 vs. –0.25 ± 0.04, Teng et al., 2010b). The following observations also suggest that dry garnet peridotite is not a suitable source for our alkaline basalts. First, alkaline basalts usually have lower SiO2 and Al2O3, and higher TiO2, Fe2O3t and CaO at a given MgO content than high-pressure experimentally-determined partial melts of dry garnet peridotite (e.g., Hirose and Kushiro, 1993; Walter, 1998; Dasgupta et al., 2007). Second, partial melting of dry garnet peridotite cannot produce the superchondritic Zr/Hf ratios observed in our basalts (Table 1) because Zr and Hf have similar partition coefficients between dry peridotite and basaltic melts (Salters et al., 2002). Third, because the bulk partition coefficients for Zr, Hf and Ti between garnet peridotite and basaltic melts are similar to those of middle REEs (Sm, Eu and Gd) (Salters et al., 2002), the negative anomalies of Zr, Hf and Ti relative to neighboring REEs cannot be explained by partial melting of dry garnet peridotite but are consistent with the features of carbonatites as shown in Fig. 4. The low δ26Mg values (as low as –0.60, Figs. 6-8) of alkaline basalts from Anfengshan and Pingmingshan, which were generated by low degrees (2~3%, Fig. 10) of partial melting of the mantle, suggest that the mantle source for our basalts has much lighter Mg isotopic compositions relative to the normal mantle (–0.25 ± 0.7, Teng et al., 2010b). This indicates that the mantle source for our basalts may be metasomatized by isotopically light carbonatitic melts, because the deeply recycled carbonates and carbonated eclogites have light Mg isotopic compositions, with δ26Mg of –2.51 to –0.53 (Wang et al., 2014). Additionally, the decrease of the total alkalis (i.e., Na2O + K2O) and TiO2 with increasing degree of melting (Fig. 7a, b) is 12

consistent with the compositional trends observed in carbonated silicate melts produced by partial melting of fertile natural peridotite KLB-1 + 1~2.5 wt.% CO2 (Dasgupta et al., 2007, 2013). This further indicates that our basalts probably are sourced from a carbonated mantle. Furthermore, the presence of carbonatitic melts in the mantle source is also suggested by Zeng et al. (2010) based on the relationship between total alkalis (Na2O + K2O) and TiO2. In the plot of total alkalis vs. TiO2 (Fig. 12), the Cenozoic basalts from eastern China fall along the trend defined by experimentally-derived melts of carbonated peridotite, implying that carbonated peridotite is probably the main source for them. As carbonation significantly lowers the solidus of mantle peridotite, carbonated mantle may melt before anything else and contribute more where the degrees of melting are low. It has been experimentally determined that the incipient melts from carbonated mantle at near solidus are carbonatitic melts (Dasgupta et al., 2007; 2013). Carbonatitic melts will evolve to carbonated silicate melts with increasing temperature and degree of melting as the dissolution of clinopyroxene and/or olivine into carbonatitic melts becomes significant, melt fraction increases and the concentration of CO2 in melts gets diluted (Dasgupta et al., 2007; 2013). Therefore, under high degrees of melting and higher temperatures, the so-called “carbonatitic fingerprints” (e.g., high Ca/Al and Zr/Hf ratios, extremely low Ti/Ti* and Hf/Hf*, and strongly negative anomalies of Zr, Hf and Ti in spidergram, Hoernle et al., 2002; Bizimis et al., 2003) will be diluted as shown in the studied basalts from Fangshan and Chongren (Figs. 3 and 4). The features that the δ 26Mg values in our basalts increase with increasing degree of melting could be attributed to the incongruent melting of the carbonated mantle. The basalts produced by low degrees of melting were contributed more from istopically light carbonatitic melts and thus show much lighter Mg isotopic compositions, as observed in basalts from Anfengshan and Pingmingshan; while the basalts produced by higher degrees of melting were contributed largely from isotopically heavy silicate melts and thus display heavy Mg isotopic compositions, as observed in basalts from Chongren. As illustrated in Fig. 13, the basalts with lower δ26Mg values generally have higher Ca/Al and Zr/Hf ratios, and lower Hf/Hf* and Ti/Ti* ratios. Because carbonatitic melts have high ratios of Ca/Al and Zr/Hf and low ratios of Hf/Hf* and Ti/Ti* (Hoernle et al., 2002; Bizimis et al., 2003), these features also imply that our basalts probably record the compositional evolution trend from carbonatitic melts with light Mg isotopic compositions to silicate melts with heavy isotopic compositions (represented by N-MORB in Fig. 13) as observed in partial melting experiments on carbonated peridotite (e.g., Dasgupta et al., 2007, 2013). Thus, the mantle source for our basalts from the SCB is a carbonated mantle with light Mg isotopic compositions that formed 13

by incorporation of isotopically light carbonatitic melts into the upper mantle. As the studied basalts are enriched in LILE and LREE and have depleted Sr-Nd isotopic compositions (e.g., Zou et al., 2000; Chen et al., 2009; Wang et al., 2011), the incorporation of carbonatitic components into the upper mantle must take place recently without a long time-integrated ingrowth of Sr-Nd isotopic systems.

5.4. Do the carbonatitic components come from the subducted oceanic slab?

The carbonatitic melts that metasomatize the upper mantle to form carbonated peridotite for generating the alkaline basalts may come from the deep mantle at great depths or the subducted slabs (Dasgupta et al., 2007; 2013). It is experimentally shown that the carbonatitic melts from the deep mantle can be generated by partial melting or redox melting of the primitive carbon-bearing peridotite at depths of greater than 200 km (Dasgupta and Hirschman, 2006; Dasgupta et al., 2013; Stagno et al., 2013), where carbon is stored chiefly in carbonates (e.g., dolomite and magnesite), graphite/diamond and carbides (e.g., Luth, 1999; Dasgupta and Hirschmann, 2010). The graphite or diamond in the deep mantle would react with silicates to form dolomite or magnesite during their ascent through the following redox reactions proposed by Eggler (1982) and Luth (1993):

2C

graphite / diamond

+ 2O2 + 2Mg 2SiO4 = Mg 2Si 2O6 + 2MgCO3 olivine

enstatite

(1)

magnesite

and

CaMgSi2O6 + 2Mg 2SiO4 + dioposide

olivine

2C

graphite / diamond

+ 2O2 = 2Mg 2Si 2O6 + CaMg(CO3 )2 enstatite

(2)

dolomite

Both reactions are involved with free O2 as oxidants, which could originate from the mantle-derived fluids through reduction of oxidized species, such as CO2, H2O and sulfates (e.g., Luth, 1993). These oxidized species are usually abundant in mantle-derived fluids, manifested by the study of fluid inclusions in diamond and minerals of spinel and garnet-peridotite xenoliths (e.g., Navon et al., 1988; Frezzotti et al., 2012). According to these reactions, the element Mg in these carbonates is mainly derived from olivine and pyroxene that have mantle-like Mg isotopic compositions in the normal mantle (e,g., Hardler et al., 2009; Yang et al., 2009; Huang et al., 2011; Liu et al., 2011; Pogge von Strandmann et al., 2011; Xiao et al., 2013). Recent experimental study shows that Mg isotope fractionation between olivine and carbonate (e.g., magnesite) is limited at temperatures of ≥ 800 oC (Δ26Mgolivine-magnesite ≤ 0.04 ± 0.04, Macris et al., 2013), suggesting that primitive carbonate 14

and olivine have similar Mg isotopic compositions under mantle temperatures. Therefore, the deep mantle-derived carbonatitic melts are inferred to have normal mantle-like Mg isotopic compositions and cannot produce the low δ26Mg values of our basalts. Sedimentary carbonates so far reported have the lightest Mg isotopic compositions, with δ26Mg of –5.54 to –0.47 (e.g., Galy et al., 2002; Young and Galy, 2004; Tipper et al., 2006; Pogge von Strandmann et al., 2008a; Higgins and Schrag, 2010; Ke et al., 2011; Pokrovsky et al., 2011; Wombacher et al., 2011). A recent study revealed differential isotopic exchange between the eclogites and carbonates (e.g., limestone to dolostone) during subduction and found that the deeply recycled carbonates and carbonated eclogites have light Mg isotopic compositions, with δ26Mg of –2.51 to –0.53 (Wang et al., 2014). Thus, partial melting of rocks containing low δ26Mg carbonates will generate low δ26Mg carbonatitic melts. Metasomatism of the mantle by isotopically light carbonatitic melts from the subducted slab could form carbonated peridotite and shift the Mg isotopic compositions of the mantle to light values (Yang et al., 2012; Xiao et al., 2013). The carbonatitic melts were probably derived from partial melting of the recycled carbonated eclogite transformed from the subducted carbonate-bearing oceanic crust during plate subduction. This can be inferred from the following observations. First, the OIB-like trace element distribution patterns (Fig. 4) and key element ratios (e.g., U/Pb, Th/Pb, Nb/U, Ce/Pb, Table 1) for most of the basalts imply the involvement of a few percent of recycled oceanic crust in their mantle source (e.g., Wang et al., 2011; Xu et al., 2012; Sakuyama et al., 2013). Second, carbonated eclogite has been reported to have light Mg isotopic compositions, with δ26Mg as low as –1.93 (Wang et al., 2012, 2014).

5.5. Geodynamic implications The light Mg isotopic compositions of <110 Ma basalts from the NCB have been attributed to the interaction of mantle peridotite with isotopically light carbonatitic melts derived from the subducted oceanic slab (Yang et al., 2012). However, the NCB suffered from three circum-craton oceanic subductions since the Paleozoic era, including the Paleo-Tethys oceanic subduction from south, the Mongolia oceanic subduction from north and the Pacific oceanic subduction from East (Windley et al., 2010). It is thus difficult to judge which oceanic subduction supplied recycled carbonated eclogite in the upper mantle of the NCB. Previous studies show that on the northern side of the NCB, southwards subduction of the Mongolia oceanic plate started in the Ordovician and ceased in the Permo-Triassic (Xiao et al., 2003), while northwards subduction of the Paleo-Tethys oceanic plate below the 15

southern margin of the NCB began in the Paleozoic and ceased in the Triassic (Li et al., 1993, 2000, 2001). On the basis of plate reconstruction, westwards subduction of the Izanaghi-Pacific plate beneath the eastern Asian continent was suggested to start as early as since early Cretaceous (Müller et al., 2008). Meanwhile, the mantle-like Mg isotopic compositions of >120 Ma basalts in the NCB (Fig. 6) suggest that the isotopically light mantle source didn’t form till 120Ma (Yang et al., 2012). Considering only the Pacific plate was subducting beneath the NCB in the Mesozoic-Cenozoic, Yang et al. (2012) further pointed out that subduction of Pacific plate plays an important role in generating the light Mg isotopic compositions of the <110 Ma basalts of the NCB. In this study, the light Mg isotopic compositions of the Cenozoic basalts from the SCB provides a convincing evidence to support that the recycled carbonated eclogite in the upper mantle of the NCB were derived from the Pacific slab, because only the Pacific slab has an influence on both blocks of North and South China. In addition, our results combined with Yang’s study also suggest that the upper mantle of eastern China may be significantly metasomatized by carbonatitic melts formed by partial melting of carbonated eclogite transformed from the subducted Pacific slab. Such carbonatitic melts might be responsible for the abrupt changes of Mg and Nd isotopic compositions between the >120 and <110 Ma continental basalts from eastern China (Yang et al., 2012). High-resolution seismic tomography revealed that the Pacific slab is subducting beneath the Japan Islands and becomes stagnant in the mantle transition zone (410–660 km) beneath eastern China, with its western edge ~2000 km away from the Japan Trench (e.g., Zhao et al., 2011). The stagnant Pacific slab might bring large amounts of carbonated eclogites into the mantle transition zone as they can survive from subduction-zone dehydration and melting at modern subduction zones (Dasgupta, 2013 and references therein). Thus, we propose that the carbonatitic melts for creating the light Mg isotopic compositions of the upper mantle of eastern China probably originate from the stagnant Pacific slab. Partial molten carbonatitic melts would metasomatize the upper mantle and lead to the formation of a carbonated mantle above the stagnant Pacific slab beneath East Asia, which have formed a big mantle wedge (e.g., Zhao et al., 2011). Carbonated peridotite with light Mg isotopic compositions existed in the upper mantle would melt first and generate the isotopically light basalts, because carbonation lowers the solidus of peridotite (e.g., Dasgupta et al., 2007, 2013).

6. Conclusions 16

High-precision major and trace element data and Mg isotopic analyses on the Cenozoic alkaline and tholeiitic basalts from the SCB, eastern China lead us to make conclusions as follows: (1)

The Mg isotopic compositions of the studied basalts are much lighter relative to the

normal mantle, with δ26Mg values ranging from –0.60 to –0.35. The possibility of isotopically light ilmenite accumulation in their mantle source for causing the light Mg isotopic compositions of our basalts can be ruled out, because (i) their relatively lower TiO2 contents (< 2.5 wt.%) compared to high-Ti lunar basalts (> 6.5 wt.%) investigated by Sedaghatpour et al. (2013) suggest no significant abundance of ilmenite in the mantle source of our basalts; and (ii) the roughly positive correlation between their δ26Mg values and Ti/Ti* as well as their constant Nb/Ta ratios irrespective of variable TiO2 contents suggests no isotopically light ilmenite accumulation in their mantle source, which would result in negative correlations between Nb/Ta and TiO2, δ26Mg and Ti/Ti*. Thus, the low δ26Mg basalts from the SCB were probably sourced from a carbonated mantle that formed by interaction of the mantle with isotopically light carbonatitic melts, and our results confirm that Mg isotope ratios can be used as a powerful tool to trace recycled carbonates. (2) The δ26Mg values of our basalts decrease with the amounts of incompatible elements (e.g., Ti, La, Nd, Nb, Th) and trace element abundance ratios (e.g., Sm/Yb, Nb/Y) that are sensitive to partial melting, suggesting that large variations in Mg isotope ratios occurred during partial melting of the mantle under high temperatures and pressures. Additionally, their Hf/Hf* and Ti/Ti* ratios increase, and Ca/Al and Zr/Hf ratios decrease with increasing degrees of partial melting. These features can be ascribed to the incongruent partial melting of the carbonated mantle. At low degrees of melting, the partial melts are contributed more from isotopically light carbonatitic melts that have high Ca/Al and Zr/Hf ratios, low Hf/Hf* and Ti/Ti* ratios, while at higher degrees of partial melting, the partial melts are contributed more from isotopically heavy silicate melts that have low Ca/Al and Zr/Hf ratios, high Hf/Hf* and Ti/Ti* ratios. Thus, the large Mg isotopic variations in our basalts represent conservative mixing of isotopically distinct materials rather than isotope fractionation at mantle pressures and temperatures. (3) The carbonatitic melts probably originate from the stagnant Pacific slab beneath East Asia, which is consistent with the results of seismic tomography (e.g., Zhao et al., 2011). Thus, our results combined with the study of Yang et al. (2012) demonstrate that the subducted Pacific slab provides the recycled carbonates and that there exists a widespread carbonated upper mantle beneath eastern China, which serves as the main source for the 17

<110Ma alkaline basalts.

Acknowledgement

This work is financially supported by grants from the National Science Foundation of China (NOs. 41090372, 41230209, 41328004 and 41273037) and the Fundamental Research Funds for the Central Universities (WK2080000068). We are grateful to Zhen-Hui Hou and Hai-Yang Liu for assistance during analyses of trace elements and to Ting Gao, Zi-Jian Li, Ze-Zhou Wang for Mg isotopic analyses. Grateful thanks are due to four anonymous reviewers whose constructive and critical reviews greatly improve this contribution. We acknowledge the executive editor Dr. Marc Norman and the associated editor Dr. Weidong Sun for their punctuality and dedication during the editing work.

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Lee T., Papanastassiou D. A. and Wasserburg G. J. (1976) Demonstration of

26

Mg excess in Allende and

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24

Figure Captions Fig. 1. Simplified geological map of eastern China that mainly consists of the SCB, the NCB, and NE China (i.e., the Xing-Meng Block). The regions where basalts were sampled for Mg isotope investigation in this study are marked as yellow triangles.

Fig. 2. Na2O + K2O vs. SiO2 diagram (Le Bas et al., 1986) for the SCB Cenozoic basalts. Fig. 3. MgO vs. other oxides diagrams for the SCB Cenozoic basalts.

Fig. 4. Chondrite-normalized REE patterns (a) and primitive mantle (PM)-normalized trace element distribution patterns (b) of the SCB Cenozoic basalts. Normalized values are from McDonough and Sun (1995), and data for N-MORB and OIB are from Sun and McDonough (1989). The average values for magnesio- and calico-carbonatites are taken from references (Hoernle et al., 2002; Bizimis et al., 2003).

Fig. 5. Ti/Ti* vs. Ca/Al (a) and Hf/Hf* (b) in the SCB Cenozoic basalts. Yellow stars represent the average ratios of Ca/Al (248.4), Ti/Ti* (0.106) and Hf/Hf* (0.016) for magnesio- and calico-carbonatites (Hoernle et al., 2002; Bizimis et al., 2003). Those ratios for N-MORB are calculated based on major and trace element data presented by Hofmann (1988). Fig. 6. (a) δ26Mg vs. δ25Mg in the SCB Cenozoic basalts and the USGS standards. It is noted that all data distribute along the terrestrial equilibrium mass fractionation line with a slope of 0.521 (Young and Galy, 2004); (b) δ26Mg vs. Mg# in the SCB Cenozoic basalts. Also shown for comparison are Mg isotopic data of fresh basaltic lavas from the NCB (Yang et al., 2012). Gray bar represents the widely accepted δ 26Mg of the normal mantle (–0.25 ± 0.07, Teng et al., 2010b). Fig. 7. Variations of δ26Mg with the abundance of incompatible elements in the SCB Cenozoic basalts. The δ26Mg of N-MORB is cited as –0.25 ± 0.07 (Teng et al., 2010b) and other values for incompatible elements are taken from Hofmann (1988). Fig. 8. δ26Mg vs. Sm/Yb (a) and Nb/Y (b) in the SCB Cenozoic basalts. The negative correlations between δ26Mg and Sm/Yb and Nb/Y, trace element abundance ratios that are 25

sensitive to partial melting, suggest that the observed δ 26Mg variation in the SCB Cenozoic basalts is caused by partial melting of a carbonated mantle (See text for details). Fig. 9. Variations of selected elements versus Nb in the SCB Cenozoic basalts

Fig. 10. La/Yb vs. Sm/Yb in the SCB Cenozoic basalts. Also shown is the batch melting curve calculated for garnet peridotite. Partition coefficients are taken from Johnson et al. (1990). The starting material are ol, 55%; opx, 20%; cpx, 15%; grt, 15%; melting reaction in garnet field (Walter, 1998): ol, 3%; opx, 3%; cpx, 70%; grt, 24%. The inverse modeling used here follows Feigenson et al. (2003) and Xu et al. (2005).

Fig. 11. TiO2 vs. Nb/Ta in the SCB Cenozoic basalts. Fig. 12. Na2O + K2O vs. TiO2 for the Cenozoic basalts in eastern China. Data sources for the Cenozoic basalts in eastern China are from the references (Zhi et al., 1990; Zou et al., 2000; Xu et al., 2005; Tang et al., 2006; Liu et al., 2008; Chen et al., 2009; Zhang et al., 2009; Zeng et al., 2010, 2011; Wang et al., 2011; Xu et al., 2012). Also shown for comparison are experimentally-derived melts from carbonated peridotite (Hirose, 1997; Dasgupta et al., 2007), carbonated pyroxenite (Gerbode and Dasgupta, 2010), pyroxenite or eclogite (Hirschmann et al., 2003; Kogiso et al., 2003; Pertermann and Hirschmann, 2003; Kogiso and Hirschmann, 2006), hornblendite (Pilet et al., 2008), and carbonated eclogite (Dasgupta et al., 2006). Fig. 13. δ26Mg vs. Ca/Al (a), Zf/Hf (b), Hf/Hf* (c) and Ti/Ti* (d) in the SCB Cenozoic basalts.

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39

Table 1. Major and trace element concentrations of the Cenozoic basalts from the South China Block. Sample a Location

13PMS1 13PMS1R 13PMS2

13PMS4

13PMS5

13PMS6

13PMS7

40.86 2.39 11.70 13.18 0.20 9.93 9.36 4.50 2.38 1.25 3.22 98.97 6.88 1.89 0.64

41.08 2.43 11.74 13.23 0.20 10.00 9.43 4.63 2.32 1.29 3.24 99.59 6.95 2.00 0.64

41.05 2.41 11.86 13.00 0.20 9.71 9.58 4.29 2.46 1.23 3.36 99.15 6.75 1.74 0.64

41.14 2.41 11.73 13.20 0.20 10.11 9.47 4.43 2.27 1.26 3.35 99.57 6.70 1.95 0.64

41.07 2.41 11.66 13.35 0.20 10.10 9.50 4.60 2.33 1.28 3.00 99.50 6.93 1.97 0.64

40.91 2.39 11.58 13.24 0.20 10.28 9.53 4.22 2.18 1.26 3.61 99.40 6.40 1.94 0.65

11.9 4.02 12.6 130 217 46.4 202 51.8 21.9 38.5 1490 32.9 335 124 1.49 672 78.7 141 15.9 65.1 14.9 4.68 13.0 1.79 8.65 1.28 2.87 0.32 1.75 0.21 6.71 6.71 5.47 11.6 3.13 0.73 39.7 25.8 45.0 8.5 18.5 49.9 0.45 0.54

9.76 5.73 14.4 136 225 49.1 217 45.1 26.1 39.3 1557 33.6 312 116 1.32 592 71.4 129 14.1 61.8 13.8 4.43 14.6 1.62 7.51 1.16 2.60 0.28 1.74 0.26 6.14 5.96 5.71 11.0 3.18 0.73 36.4 22.6 41.0 7.9 19.4 50.9 0.43 0.53

11.7 4.16 12.5 125 198 46.1 197 54.5 21.9 45.9 1583 33.6 336 126 1.49 697 79.1 127 16.0 65.5 15.0 4.67 13.1 1.80 8.66 1.26 2.91 0.32 1.79 0.22 6.65 6.68 5.58 11.7 3.09 0.74 40.8 22.8 44.2 8.4 18.9 50.5 0.45 0.53

11.7 4.01 12.7 130 216 46.4 206 51.6 22.6 37.9 1348 32.7 328 119 1.47 648 75.9 137 15.6 63.9 14.6 4.57 12.8 1.75 8.38 1.22 2.78 0.31 1.70 0.21 6.53 6.44 5.23 11.3 3.14 0.74 37.9 26.2 44.6 8.6 18.5 50.2 0.46 0.54

12.2 4.15 13.0 133 223 48.3 211 53.5 22.9 38.9 1417 33.3 333 120 1.50 634 75.2 127 15.4 62.9 14.3 4.46 12.5 1.70 8.24 1.20 2.73 0.30 1.67 0.21 6.45 6.28 5.34 11.1 3.09 0.74 38.9 23.8 45.0 8.6 19.1 51.6 0.47 0.54

12.2 4.20 13.2 139 232 48.7 220 53.7 23.7 41.7 1420 33.3 334 119 1.36 634 75.5 135 15.3 62.6 14.2 4.49 12.6 1.73 8.28 1.20 2.75 0.31 1.69 0.21 6.43 6.27 5.31 11.2 3.06 0.75 38.9 25.4 44.7 8.4 19.0 51.9 0.47 0.54

Pingmingshan

Major element (wt.%) SiO2 41.02 TiO2 2.43 Al2O3 11.71 Fe2O3 13.34 MnO 0.20 MgO 10.03 CaO 9.52 Na2O 4.73 K2O 2.23 P2O5 1.25 LOI 3.05 Total 99.51 Na2O+K2O 6.96 Na2O/K2O 2.12 Mg# b 0.64 Trace element (ppm) Li Be Sc V Cr Co Ni Cu Ga Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U Ca/Al c Nb/U Ce/Pb La/Yb Sm/Yb Nb/Ta Zr/Hf Ti/Ti* d Hf/Hf* e

13PMS3

10.4 3.33 12.9 119 197 43.3 193 56.1 20.7 34.7 1211 30.2 307 112 1.49 642 77.6 139 15.6 64.3 14.7 4.58 13.1 1.80 8.67 1.27 2.85 0.32 1.75 0.22 6.59 6.46 5.86 11.1 2.81 0.74 39.9 23.7 44.3 8.4 17.3 46.6 0.46 0.54

11.2 3.96 12.8 122 203 44.9 197 57.9 22.3 35.9 1275 31.8 316 113 1.41 613 73.0 131 14.8 60.5 13.7 4.28 12.0 1.65 8.00 1.17 2.64 0.29 1.60 0.20 6.14 5.99 5.57 10.8 2.73 41.5 23.5 45.6 8.6 18.9 51.5

40

Sample

13PMS8 Pingmingshan

Major element (wt.%) SiO2 41.17 TiO2 2.64 Al2O3 12.48 Fe2O3 14.58 MnO 0.22 MgO 8.03 CaO 9.24 Na2O 6.44 K2O 2.46 P2O5 1.13 LOI 1.39 Total 99.78 Na2O+K2O 8.90 Na2O/K2O 2.62 Mg# 0.56 Trace element (ppm) Li 13.5 Be 5.27 Sc 11.4 V 119 Cr 104 Co 44.1 Ni 113 Cu 55.2 Ga 27.6 Rb 29.7 Sr 1468 Y 36.8 Zr 363 Nb 154 Cs 1.60 Ba 585 La 82.2 Ce 125 Pr 17.1 Nd 71.4 Sm 16.3 Eu 5.10 Gd 14.2 Tb 1.95 Dy 9.36 Ho 1.35 Er 3.01 Tm 0.33 Yb 1.75 Lu 0.21 Hf 8.30 Ta 8.56 Pb 5.66 Th 12.2 U 4.12 Ca/Al 0.67 Nb/U 37.4 Ce/Pb 22.1 La/Yb 47.0 Sm/Yb 9.3 Nb/Ta 18.0 Zr/Hf 43.7 Ti/Ti* 0.46 Hf/Hf* 0.61

13AFS1

13AFS2

13AFS3

13AFS4

13AFS9

13AFS10

Anfengshan

10FS6 Fangshan

44.46 2.36 12.05 12.48 0.20 7.26 8.77 6.19 2.87 1.43 1.63 99.70 9.06 2.16 0.58

44.41 2.37 12.04 12.50 0.20 7.24 8.74 6.22 2.78 1.42 1.60 99.52 9.00 2.24 0.58

42.41 2.49 11.89 13.06 0.21 8.44 9.08 5.59 2.15 1.27 2.90 99.49 7.74 2.60 0.60

42.36 2.49 11.86 13.04 0.21 8.36 9.08 5.61 2.17 1.27 2.88 99.33 7.78 2.59 0.60

40.86 2.34 11.73 12.57 0.20 10.54 9.61 3.73 1.88 1.35 4.66 99.47 5.61 1.98 0.66

40.65 2.32 11.51 12.55 0.20 10.13 9.54 4.20 2.01 1.44 4.11 98.66 6.21 2.09 0.66

46.56 2.12 13.89 11.35 0.15 9.72 7.43 3.66 2.60 0.54 1.60 99.62 6.26 1.41 0.67

13.8 5.32 10.4 112 107 37.6 97.9 43.0 25.7 42.0 1429 42.4 372 138 1.91 728 92.6 140 19.4 79.3 17.8 5.08 15.3 2.14 10.30 1.54 3.51 0.42 2.35 0.31 8.50 7.19 7.72 14.4 3.10 0.66 44.6 18.1 39.4 7.6 19.2 43.8 0.38 0.57

17.8 9.03 12.8 120 114 40.0 108 36.2 29.4 39.3 1632 43.7 423 139 1.86 701 92.1 169 18.5 82.1 17.7 5.09 19.4 2.10 9.57 1.49 3.35 0.39 2.45 0.35 8.11 7.08 7.55 13.8 3.27 0.66 42.4 22.4 37.6 7.2 19.6 52.1 0.32 0.54

14.3 5.59 11.0 164 111 40.2 106 47.5 27.8 37.4 1596 42.4 385 148 1.66 699 92.6 141 19.5 79.7 18.3 5.39 15.9 2.18 10.40 1.51 3.45 0.39 2.19 0.28 8.47 7.76 7.28 13.9 2.87 0.70 51.6 19.4 42.3 8.4 19.1 45.5 0.38 0.56

14.4 5.33 10.9 209 113 39.6 103 47.5 26.9 40.0 1566 41.3 371 141 1.81 675 92.4 140 19.1 78.8 18.1 5.35 15.6 2.14 10.30 1.49 3.39 0.38 2.15 0.27 8.36 7.53 7.38 13.6 2.89 0.70 48.9 19.0 43.0 8.4 18.7 44.4 0.39 0.56

14.7 5.26 14.1 131 186 44.6 168 60.9 24.0 32.0 1518 39.5 372 130 1.28 633 84.3 125 17.3 71.0 16.3 5.01 14.0 1.94 9.30 1.36 3.08 0.35 1.97 0.25 6.99 6.48 6.16 13.1 3.45 0.75 37.7 20.3 42.8 8.3 20.1 53.2 0.41 0.52

16.6 5.15 14.3 132 172 44.7 160 51.5 24.4 50.5 1823 40.5 343 133 1.66 701 87.4 132 18.0 73.3 17.0 5.23 14.8 2.02 9.66 1.41 3.20 0.36 2.04 0.26 7.20 6.70 6.27 13.3 3.34 0.75 39.9 21.1 42.8 8.3 19.9 47.6 0.38 0.51

9.28 2.29 15.9 145 323 49.2 296 54.8 21.3 45.4 585 20.4 197 56.2 0.82 437 28.8 54.6 6.25 25.1 5.76 1.92 5.39 0.82 4.34 0.71 1.78 0.23 1.39 0.19 4.04 3.29 3.09 4.0 1.18 0.49 47.8 17.7 20.7 4.1 17.1 48.8 0.98 0.85

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Sample

10FS8

10FS9

Fangshan Major element (wt.%) SiO2 46.66 TiO2 2.10 Al2O3 13.72 Fe2O3 11.12 MnO 0.15 MgO 10.08 CaO 7.20 Na2O 3.90 K2O 1.68 P2O5 0.53 LOI 2.23 Total 99.37 Na2O+K2O 5.58 Na2O/K2O 2.32 Mg# 0.68 Trace element (ppm) Li 9.13 Be 2.12 Sc 15.6 V 130 Cr 347 Co 48.4 Ni 307 Cu 60.9 Ga 21.0 Rb 32.6 Sr 608 Y 19.9 Zr 191 Nb 55.0 Cs 2.10 Ba 421 La 27.6 Ce 52.9 Pr 6.03 Nd 24.3 Sm 5.53 Eu 1.83 Gd 5.08 Tb 0.77 Dy 4.11 Ho 0.68 Er 1.68 Tm 0.22 Yb 1.30 Lu 0.18 Hf 3.92 Ta 3.20 Pb 2.69 Th 3.9 U 1.15 Ca/Al 0.48 Nb/U 48.0 Ce/Pb 19.7 La/Yb 21.2 Sm/Yb 4.3 Nb/Ta 17.2 Zr/Hf 48.7 Ti/Ti* 1.03 Hf/Hf* 0.85

10FS10

10FS11

Fangshan

10CR1

10CR2

Chongren

10LYSK11

10LYSK13

Longyou

46.73 2.12 13.84 11.22 0.15 9.99 7.29 3.94 1.61 0.53 2.31 99.73 5.55 2.45 0.68

46.48 2.10 13.72 11.13 0.15 9.99 7.27 3.71 1.74 0.53 2.48 99.30 5.45 2.13 0.68

46.44 2.13 13.88 11.33 0.16 10.11 7.56 3.67 2.12 0.53 1.89 99.82 5.79 1.73 0.68

51.10 2.29 11.87 13.08 0.18 6.91 6.07 3.27 1.46 0.58 3.06 99.87 4.73 2.24 0.55

51.13 2.14 13.47 12.23 0.17 7.49 8.54 3.22 1.04 0.34 -0.08 99.69 4.26 3.10 0.59

43.4 1.81 10.1 10.78 0.16 13.05 9.02 2.99 2.07 0.91 4.94 99.23 5.06 1.44 0.74

42.8 2.14 11.4 12.2 0.18 12 9.87 4.01 1.44 1.11 2.15 99.30 5.45 2.78 0.70

9.38 2.11 15.8 145 355 50.1 304 62.4 21.2 32.9 609 19.8 192 55.6 1.73 425 27.9 53.3 6.07 24.4 5.56 1.84 5.15 0.78 4.15 0.68 1.69 0.22 1.33 0.18 3.93 3.21 2.69 4.0 1.15 0.48 48.5 19.8 21.0 4.2 17.3 48.9 1.02 0.85

9.49 2.25 15.7 144 351 48.7 303 63.4 20.9 32.5 630 19.9 190 55.1 1.51 427 28.1 53.6 6.09 24.5 5.67 1.84 5.20 0.78 4.13 0.68 1.69 0.22 1.35 0.18 3.99 3.20 2.84 4.0 1.16 0.48 47.7 18.9 20.8 4.2 17.2 47.6 1.00 0.85

8.55 2.05 15.7 140 397 47.5 303 58.3 19.3 41.6 557 19.3 184 53.2 1.10 455 30.4 58.3 6.66 27.2 6.23 2.08 5.74 0.87 4.69 0.75 1.88 0.25 1.46 0.20 4.06 3.23 3.09 3.8 1.09 0.50 49.0 18.9 20.8 4.3 16.5 45.3 0.92 0.79

7.10 1.85 13.4 135 108 40.5 170 112 22.6 33.9 298 33.4 221 30.0 1.03 544 27.8 55.7 7.01 31.9 9.15 2.76 8.78 1.43 7.72 1.30 3.24 0.43 2.61 0.35 5.47 1.62 4.10 3.7 0.78 0.47 38.7 13.6 10.7 3.5 18.5 40.4 0.65 0.81

6.47 0.99 19.0 161 168 45.1 157 79.7 18.7 20.6 344 20.7 125 20.2 0.41 320 15.2 31.0 3.89 18.2 5.35 1.89 5.48 0.91 4.99 0.86 2.11 0.28 1.73 0.24 3.36 1.23 2.46 2.1 0.49 0.58 41.6 12.6 8.8 3.1 16.4 37.2 0.99 0.86

16.9 2.07 19.3 143 504 55.0 451 55.7 17.3 64.4 974 25.1 219 88.8 1.1 648 57.1 106 11.7 48.0 9.43 2.82 8.13 1.08 5.41 0.87 2.15 0.27 1.55 0.21 4.56 4.65 4.34 9.0 1.98 0.81 44.8 24.4 36.8 6.1 19.1 48.0 0.53 0.54

9.47 2.35 19.4 152 375 54.0 349 56.5 18.9 38.3 1194 29.2 244 107 1.06 915 68.9 126 14.2 57.8 11.4 3.44 9.55 1.27 6.28 1.02 2.46 0.3 1.74 0.23 4.99 5.54 3.75 10.0 2.18 0.79 49.1 33.6 39.6 6.6 19.3 48.9 0.47 0.49

a

13PMS1R is the replicated analysis of sample 13PMS1 for trace element concentrations. Mg# = Mg/(Mg + 0.85Fetot). c Ca/Al = ([CaO]wt.%/56)/(2*[Al2O3]wt.%/102) b

42

d e

Ti/Ti* = TiN/(NdN‒0.055 × SmN0.333 × GdN0.722) Hf/Hf* = HfN/(SmN × NdN)0.5

43

Table 2. Magnesium isotopic compositions of the Cenozoic basalts from the South China Block. Sample

δ26Mga

2SDb

δ25Mg

2SDb

N

Δ25Mg'c

13PMS1 13PMS2 13PMS3 Replicated 13PMS6 13PMS8 13AFS1 Replicate 13AFS2 13AFS3 Replicate 13AFS4 10FS6 Replicate 10FS8 10FS9 10FS10 10FS11 10CR1 10CR2 10LYSK13 BHVO‒2 Replicate AGV‒2 Replicate GSP‒2

‒0.52 ‒0.53 ‒0.52 ‒0.52 ‒0.52 ‒0.60 ‒0.52 ‒0.51 ‒0.52 ‒0.59 ‒0.58 ‒0.56 ‒0.41 ‒0.42 ‒0.42 ‒0.43 ‒0.42 ‒0.40 ‒0.37 ‒0.35 ‒0.41 ‒0.26 ‒0.24 ‒0.16 ‒0.15 0.03

0.07 0.04 0.05 0.05 0.05 0.05 0.04 0.08 0.08 0.04 0.05 0.04 0.07 0.06 0.05 0.05 0.04 0.03 0.05 0.06 0.05 0.04 0.06 0.03 0.02 0.09

‒0.25 ‒0.26 ‒0.24 ‒0.24 ‒0.26 ‒0.31 ‒0.26 ‒0.26 ‒0.25 ‒0.29 ‒0.29 ‒0.28 ‒0.22 ‒0.21 ‒0.20 ‒0.23 ‒0.22 ‒0.20 ‒0.19 ‒0.18 ‒0.21 ‒0.14 ‒0.12 ‒0.10 ‒0.06 0.00

0.09 0.03 0.03 0.03 0.03 0.05 0.02 0.05 0.07 0.04 0.04 0.02 0.03 0.02 0.03 0.05 0.04 0.05 0.04 0.05 0.05 0.06 0.02 0.06 0.04 0.06

4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4 4

0.03 0.02 0.03 0.03 0.01 0.01 0.01 0.00 0.02 0.02 0.02 0.02 0.00 0.01 0.01 ‒0.01 ‒0.01 0.00 0.00 0.00 0.01 0.00 0.01 ‒0.02 0.01 ‒0.01

a X

δ Mg = {(XMg/24Mg)sample/(XMg/24Mg)DSM3 ‒ 1) × 1000, where X = 25 or 26 and DSM3 is Mg solution made from pure Mg metal (Galy et al., 2003). b

2SD indicates twice the standard deviation of the population of 4 repeat measurements of a sample solution. c

Δ25Mg´ = δ25Mg´ – 0.521δ26Mg´, where δXMg´ = 1000 × ln[(δXMg + 1000)/1000] (Young and Galy, 2004). It is reported largely as a quality control on the data, with values that should be close to zero. d

Replicate denotes repeating sample dissolution, column chemistry and instrumental analysis.

44