A Kungurian Oceanic Upwelling on Yangtze Platform: Evidenced by
Transcription
A Kungurian Oceanic Upwelling on Yangtze Platform: Evidenced by
Journal of Earth Science, Vol. 26, No. 2, p. 211–218, April 2015 Printed in China DOI: 10.1007/s12583-015-0533-z ISSN 1674-487X A Kungurian Oceanic Upwelling on Yangtze Platform: Evidenced by δ13Corg and Authigenic Silica in the Lower Chihsia Formation of Enshi Section in South China Hao Yu1, Hengye Wei*2 1. Key Laboratory of Orogenic Belts and Crustal Evolution, MOE, Peking University, Beijing 100871, China 2. College of Earth Sciences, East China Institute of Technology, Nanchang 330013, China ABSTRACT: The Late Paleozoic Ice Age across Carboniferous and Permian had a significant impact on the Kungurian (Upper Cisuralian series of Permian) Chihsia Formation in South China. This resulted in a unique interval with features such as the lack of reef in Chihsian limestone, widespread stinkstone and nodular/bedded chert. The Chihsia limestone (Kungurian stage) deposited during a time of cooling was resulted from oceanic upwelling. Here we present evidence for this upwelling using several geochemical analyses: bulk organic carbon isotope, biomarker molecular geochemical data, and authigenic silica of the stinkstone member in the lower Chihsia Formation of the Kuangurian stage from the Enshi Section in western Hubei Province, South China. The lower part of the stinkstone member shows a rapid organic carbon isotope excursion with a -3‰ shift triggered by the upwelling of 13 C-depleted bottom water. The concurrent rapid increasing of authigenic silica content resulted from the enhanced supply of dissolved silica in the upwelling water mass. This upwelling at the Enshi Section also led to relative high TOC content, accounting for the widespread stinkstone in the lower Chihsia Formation during the Kungurian stage in Permian. KEY WORDS: Chihsia Formation, Enshi Section, organic carbon isotope, authigenic silica, upwelling. 0 INTRODUCTION The Late Paleozoic Ice Age (LPIA) lasted from the Mid-Carboniferous (ca. 327 Ma) to the early Late Permian (ca. 260 Ma) (Fielding et al., 2008) and is considered to have an important impact on the Phanerozoic Earth’s climate system (Frakes et al., 1992). Changes in vegetation during this period led to an icehouse climate state (Gastaldo et al., 1996). Fielding et al. (2008) recognized eight discrete glacial intervals, termed glaciations, during the LPIA. These glaciations are C1 to C4 in the Carboniferous and P1 to P4 in Permian, where the P3 glaciation was located in the Middle and Late Kungurian stage in Early Permian. Mii et al. (2012) suggested that the paleoclimate fluctuated between warm and cool from Late Sakmarian to Early Kungurian and that the Early Kungurian and Middle Artinskian were associated with a weakened latitudinal temperature gradient. Therefore, during Kungurian, the interglacial to glacial transition interval should indicate oceanic upwelling when the pole-to-equator temperature gradient was enhanced (e.g., Beauchamp and Baud, 2002). Upwelling during the LPIA was inferred from the ocean simulation in the Middle Permian (Winguth et al., 2002), Late Permian (Schoepfer et al., 2013; Kiehl and Shields, 2005) and *Corresponding author: [email protected]; [email protected] © China University of Geosciences and Springer-Verlag Berlin Heidelberg 2015 Manuscript received June 18, 2014. Manuscript accepted January 15, 2015. the whole Late Paleozoic (Montañez and Poulsen, 2013) along the eastern Panthalassic Ocean and on the lee side of the South China Block. The sedimentary feature such as the glendonites in eastern Australian during Mid–Late Permian (Jones et al., 2006) and the trace elemental analysis of brachiopod in the tropical region also suggested the Late Paleozoic upwelling (Powell et al., 2009). The Pangean phosphorites exhibited a record of Permian upwelling (Trappe, 1994). The Kungurian upwelling was inferred by the enrichment in minerals such as widespread sepiolite (Yan et al., 2005), lack of reefs in the lower Chihsian Formation (Shi and Grunt, 2000) and the associated chert nodules (Liu and Yan, 2007; Wang and Jin, 1998; Lu and Qu, 1989). The reducing sediments in the lower Chihsia Formation (Wei et al., 2012; Lu and Qu, 1989) and the biogeographic distribution of brachiopods recorded the cool-water upwelling systems in the Kungurian Chihsia Formation of South China (Shi and Grunt, 2000; Shi, 1995). However, the Kungurian oceanic upwelling research still needs additional geochemical evidences. Here, we present bulk organic carbon isotope (δ13Corg) and authigenic silica (SiO2(auth)) data constrained by molecular geochemical data in the limestones of the Chihsia Formation at the Enshi Section in South China to show the evidence of this Kungurian upwelling. 1 GEOLOGICAL SETTING The Enshi Section, the focus of this study, is located at the Tanjiaba Village, 5 km south of Enshi City, western Hubei Province in South China. The Enshi area became part of the intrashelf basin during the Middle and Late Permian (Wei and Yu, H., Wei, H., Y., 2015. A Kungurian Oceanic Upwelling on Yangtze Platform Evidenced by δ13Corg and Authigenic Silica in the Lower Chihsia Formation of Enshi Section in South China. Journal of Earth Science, 26(2): 211–218. doi: 10.1007/s12583-015-0533-z 212 Chen, 2011; Feng et al., 1997), which is equivalent to the Xiakou-Lichuan bay of Yin et al. (2014). This intrashelf basin is central-north of the South China Block, but was to the paleowest of the South China Block during Permian (Algeo et al., 2013), and thus was probably influenced by an eastern boundary current that was part of a circulation gyre within the Paleotethys Ocean (Kutzbach and Guetter, 1990). According to the paleomagnetic study (Ma and Zhang, 1986), the South China Block was located at 2.4°N during the time that the Chihsia Formation was deposited, suggesting a low-latitude tropical climate. The Chihsia Formation is widely exposed across the South China Block. The fusulinid and conodont biostratigraphic study in the Nanpanjiang Basin (Shen et al., 2007) suggest a latest Artinskian through the entire Kungurian stage for the Chihsia Formation in South China, indicating Early Permian, i.e., Cisuralian Epoch. However, the Chihsia Formation at Enshi Section unconformably overlies on Carboniferous karst limestones (Fig. 1). The lowermost of Chihsia Formation consists of 5-m-thick ferrallitic claystones resulted from weathering and a Hao Yu and Hengye Wei 2-m-thick coal bed (Fig. 1) in ascending order. Above these two siliciclastic successions also called Liangshan Formation (e.g., Tong and Shi, 2000), is a >120-m-thick Chihsia Formation limestone succession. We sampled the lower part of this limestone succession, in total ~20 m thick. This sampled interval, also called stinkstone (Lu and Qu, 1989), is composed of coarsely laminated marlstones or calcareous shale intercalated by thin-bed limestone or dolostone, locally bearing the black nodular chert in the marlstones/shales (Fig. 1). Grey thick-bedded limestones were developed at the base and top of these marlstones/shales interval (Fig. 1). The so-called stinkstones smell like the bituminous odor, suggesting high organic carbon content. The unconformity between Carboniferous karst limestone and lower Chihsian claystones represents the widespread Early Permian uplift and erosion for most of South China (Tong and Shi, 2000). Therefore, the sampled stinkstone member in the Chihsia Formation is reasonably Early–Middle Kungurian stage in age. Figure 1. The lithology and geochemical profiles of bulk organic carbon isotope (δ13Corg), authigenic silica (SiO2(auth)) and total organic carbon (TOC) in the lower Chihsia Formation at the Enshi Section, western Hubei Province, South China. Note that the TOC data is from Wei et al. (2012). A Kungurian Oceanic Upwelling on Yangtze Platform Evidenced by δ13Corg and Authigenic Silica 2 METHODS A total of 31 samples were sampled in this 20-m-thick study interval, with an average sampling-interval of 0.65 m. The Enshi Section was a new road-cut section, and thus the samples were very fresh. We clean the samples using distilled water, then dried them and powdered to smaller than 200 mesh. Sample splits (0.3 to 5 g) for bulk δ13Corg analysis were treated with 6 N HCl for 24 h to remove carbonate. The solution was then retreated with excess 6 N HCl and allowed to sit for 6 h to ensure there was no remaining carbonate. The decalcified samples (30–100 mg)+CuO wire (1 g) were added to a quartz tube and combusted at 500 °C for 1 h and 850 °C for 3 h. The carbon isotope ratio of the generated CO2 was measured in a Finnigan MAT-252 mass spectrometer. The isotopic ratio is reported in standard δ notation relative to the Vienna Peedee Belemnite (VPDB) standard. Analytical precision is better than 0.1‰. Sample splits (0.5 g) for the major elements analyses were analyzed on fused glass pellets using a Phillips PW 1500 X-ray fluorescence spectrometer. The precision of the major elements data is better than 3%. The authigenic fraction of element X was calculated as [X]–[Al]×[X/Al]detrital, where the detrital X/Al ratio was based on average upper crustal concentrations (McLennan, 2001). Excess silica (SiO2(xs)) was calculated as SiO2(total)–SiO2(illite)– SiO2(chlorite)=SiO2(total)–(m×K2O×2.36)–((Al2O3–m×K2O)×1.18) (c.f., Shen et al., 2013), on the assumption that most siliciclastic silicon is present as the clay mineral illite and chlorite (e.g., Hadjira et al., 2011). Where m is the slope of the Al2O3-K2O regression, the coefficients 2.36 and 1.18 represent the weight ratios of SiO2/(0.5×Al2O3) in clay minerals of stoichiometric composition having TO and TOT structures (i.e., illite and chlorite), respectively. Eleven samples were prepared for analysis of saturated hydrocarbon compounds. Powdered samples (~120 g) were Soxhlet extracted using chloroform for approximately 72 h. Asphaltenes were removed from the chloroform extracts by precipitation with n-hexane followed by filtration. The de-asphalted extracts were then separated into saturated, aromatic fractions and non-hydrocarbons by column chromatography, using hexane, benzene and methanol, respectively. For gas chromatography-mass spectrometry analyses, the saturated hydrocarbon fractions were performed using an Agilent 5973N mass spectrometer equipped with a HP 6890 gas chromatograph at the Research Institute of Petroleum Exploration and Development, China National Petroleum Corporation. The silica capillary column used was 60 m×0.25 mm in size, with 0.25 μm film in thickness. The sample was injected with an injection temperature of 300 °C. Helium was used as the carrier gas at 1 mL/min. The oven temperature was initially programmed at 100 °C for 5 min, then was programmed to increase from 100 to 220 °C at 4 °C/min. Afterwards, it was programmed to increase from 220 to 320 °C at 2 °C/min and to remain at the highest temperature for 20 min. For GC-MS analysis, the instrument was operated routinely in multiple ion detection mode (MID) with a mass scan range of 50–560 m/z. The ion source was operated in the electron impact (EI) mode at an electron energy of 70 eV and emission current of 200 μA. 213 3 RESULTS AND DISCUSSION 3.1 Upwelling Evidenced from the δ13Corg The bulk δ13Corg values range from -29.04‰ to -26.22‰, with an average of -28.17‰ (Table 1 and Fig. 1). It shows a gradual negative excursion from ~-26‰ to -28.80‰ in the lower part of stinkstones member, a persistent low δ13Corg of -29‰ interrupted by several episodes of heavy δ13Corg of -27.8‰ in the middle part of stinkstones member, and a rapid positive excursion to -26.7‰ in the upper part of stinkstones member (Fig. 1). This suggests a negative δ13Corg excursion event in the stinkstone member. Diagenetic processes can affect the δ13Corg values. Thermal maturation of organic matter decreases the total organic carbon (TOC) composition of rocks and tends to shift residual TOC to more 13C-rich values (Hayes et al., 1999; Popp et al., 1997). However, this thermal process would not affect the δ13Corg trends (Des Marais et al., 1992) because the study interval has similar thermal maturation level. Thus the negative excursion and variable changes of δ13Corg in this study (Fig. 1) rule out the thermal maturation change of organic matter. Migration of hydrocarbons or contamination by detrital δ13Corg from rock weathering could also affect the δ13Corg values (Meyer et al., 2013). Examination of microfacies by thin sections did not observe the migration of hydrocarbons in the 13 C-rich carbonate rock both at the base and top of the stinkstone member, and thus also rule out the migration effect. The fresh samples and careful treatment in the lab make sure that contamination of modern organic carbon was minimized and thus can not account for the large changes of δ13Corg in this study. However, the crossplot between TOC vs. δ13Corg (Fig. 2) shows a negative correlation (R2=0.7). Since there is no diagenetic effect, this negative relationship represents an environmental signal, instead of diagenetic signal. This negative correlation, e.g., the lower the TOC, the heavier the δ13Corg, can be due to the proportion between marine organic matter and terrestrial organic matter during the depositional period (Meyers, 1997; Whiticar, 1996) and/or the oceanic conditions changes such as the upwelling of anoxic alkalinity-charged, 13 C-depleted deepwaters (Werne and Hollander, 2004; Kaufman et al., 1997), volcanic CO2 input into the ocean (Korte and Kozur, 2010; Hansen, 2006; Grard et al., 2005; Berner, 2002) and methane release from the seafloor (Korte et al., 2010; Svensen et al., 2009, 2004; Retallack and Jahren, 2008). Large greenhouse gases input from the volcanism and methane hydrate seem unlikely even though there was a large scale extrusion of basaltic lavas in north-western Europe during the Carboniferous–Permian transition (Heeremans et al., 1996; Olaussen et al., 1994). The possibility that volcanism and methane impact on the 3‰-magnitude negative excursion of δ13Corg in the Chihsian Formation, could be low. Biomarker analyses may be a useful tool for identification of organic matter origin in the sediments and enable a better understanding of bulk organic carbon isotope (Fenton et al., 2007; Schwab and Spangenberg, 2004; Meyers, 1997). Our n-alkanes distribution of saturated hydrocarbon fraction (Fig. 3) ranges from n-C15 to n-C29 with the peak of n-C17, suggesting the main contribution of algae and bacteria (Schwab and Hao Yu and Hengye Wei 214 Table 1 Sample ES18- The major elements and bulk organic carbon isotope data in the lower Chihsia Formation at the Enshi Section, western Hubei Province, South China δ13Corg SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI SiO2(auth) (m) (‰) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) 10.56 -26.22 1.68 0.03 0.66 0.44 0.07 0.62 51.36 0.1 0.08 0.25 41.69 0.00 2.95 0.03 0.56 0.52 0.02 1.51 51.39 0 0.01 0.05 41.48 2.13 Depth ES18 11.31 -27.01 ES19 11.96 -27.70 ES20 12.11 -27.55 29.26 0.14 2.69 1.22 0.02 6.36 31.59 0.06 0.1 0.05 27.92 24.40 ES21 12.84 -27.77 28.28 0.16 3.19 1.34 0.01 8.16 30.27 0.03 0.16 0.04 27.72 21.89 ES22 13.22 -27.71 4.33 0.05 0.89 0.54 0.01 2.35 50.56 0.01 0.02 0.31 40.26 2.95 ES23 13.92 -28.32 40.41 0.12 2.24 0.91 0.01 9.14 23.18 0.05 0.15 0.02 23.17 35.31 ES24 14.32 -28.57 46.71 0.1 1.76 0.98 0.01 18.08 14.03 0.08 0.14 0.01 17.69 42.34 ES25 14.52 -28.69 47.35 0.09 1.61 0.82 0.01 18.13 12.71 0.12 0.13 0.01 16.92 43.32 15.03 -28.57 43.37 0.08 1.5 0.6 0.01 12.09 20.85 0.05 0.09 0.02 21.09 40.12 ES26 + ES27 15.12 ES28 15.73 -27.31 ES29 16.5 -28.60 41.4 0.07 1.26 0.61 0.01 16.26 19.43 0.1 0.08 0.02 20.40 38.60 ES30 17.55 -28.42 55.54 0.16 3.15 1.19 0.01 20.85 6.81 0.09 0.24 0.02 11.96 47.89 ES30+ 18.2 -28.76 30.51 0.11 2.08 0.79 0.01 10.21 28.78 0.07 0.12 0.02 26.85 26.09 47.43 0.05 1.05 0.29 <0.01 20.11 13.53 0.13 0.08 0.01 15.66 44.88 39.56 0.04 0.56 0.27 <0.01 16.37 21.04 0.16 0.05 0.01 21.41 38.08 ES31+ 19.12 -28.70 ES31 19.49 -28.38 ES32 20.15 -28.74 ES33 20.69 -27.78 ES34- 21.19 -29.04 ES34 21.64 -28.94 ES35 22.37 -27.97 ES36 23.02 -28.47 30.97 0.03 0.54 0.26 <0.01 12.72 28.77 0.05 0.03 0.02 26.12 29.84 ES37 23.77 -28.91 33.27 0.02 0.39 0.14 <0.01 15.6 25.5 0.06 0.03 0.01 24.40 32.32 ES37+ 24.52 -29.03 ES38 25.21 -28.66 ES39 25.83 -28.88 18.92 0.02 0.37 0.15 0.01 7.92 38.74 0.01 0.02 0.01 33.27 18.16 ES40 26.95 -28.83 31.33 0.03 0.47 0.19 <0.01 14.62 27.13 0.05 0.03 0.01 25.57 30.28 ES41 27.15 -27.87 9.351 0.02 0.33 0.17 0.01 2.42 47.76 0.05 0.01 0.01 39.21 8.80 ES42 28.02 -28.01 27.58 0.02 0.26 0.15 <0.01 13.29 30.76 0.02 0.01 0.01 27.46 27.11 LOI. Loss on ignition. Figure 2. The crossplot of TOC vs. δ13Corg in the lower Chihsia Formation at the Enshi Section, western Hubei Province, South China. Spangenberg, 2004; Hunt, 1996). However, there is one sample (ES20) which shows double peaks of n-alkanes distribution at n-C17 and n-C25 (Fig. 3). Even this, the δ13Corg of this sample (ES20) only show a little difference (<0.22‰) from the overlying two samples (ES21 and ES22, Figs. 1 and 3) which show single peak at n-C17. Some samples (ES37, ES28) showing relative abundant from n-C19 to n-C25 but still having a n-C17-peak display no difference of δ13Corg compared to their surrounding samples (Figs. 1 and 3). The 13C-depleted samples and 13C-rich samples in the lower part of stinkstones member (Fig. 1) have very similar pattern of n-alkanes distribution (Fig. 3). This suggests that the organic matter sources only had a small change and cannot account for the large-scale negative excursion of δ13Corg in this study. Enhanced upwelling carrying the reducing water masses A Kungurian Oceanic Upwelling on Yangtze Platform Evidenced by δ13Corg and Authigenic Silica 215 Figure 3. The n-alkanes distribution of saturated hydrocarbon fraction (m/z=218) in the lower Chihsia Formation at the Enshi Section, western Hubei Province, South China. with 13C-depleted dissolved inorganic carbon (DIC) can make the surface water DIC depleted in 13C (Werne and Hollander, 2004; Kaufman et al., 1997), and thus result in the 13C depletion of organic matter via photosynthesis (Walker et al., 2014). The Kungurian stage experienced a change from an interglacial to a glacial interval (Mii et al., 2012; Fielding et al., 2008), consistent with the coal bed and the stinkstone member, respectively since the coal bed represents the warm and humid climates (Hasiotis and Honey, 2000; Bohacs and Suter, 1997; Cecil, 1990). The transition to glaciation during the stinkstone member deposition caused enhanced oceanic upwelling and triggered the large-scale negative excursion of δ13Corg in this study. 3.2 Upwelling Evidenced from the SiO2(auth) Low Al2O3 contents (less than 3.2wt.%, average=1.7%, Table 1), combined with the high SiO2 contents (Table 1), indicate that most of the SiO2 may not be derived from continental detritus. After eliminating the continental silica, the excess silica may be mainly derived from the chert and sepiolite which contains Si and Mg elements (e.g., Yan et al., 2005). Observation during our fieldwork shows that chert bands were common Hao Yu and Hengye Wei 216 in the stinkstone member of the Chihsia Formation (Fig. 4), indicating that the diagenetic chert contribute to the excess silica. Several thin-bed dolomite layers were developed in the stinkstone member (Fig. 5). Therefore, the relatively high Mg contents in the stinkstone member (Table 1) might be resulted from these dolomite minerals. In addition, the sepiolite also contains Mg element (Yan et al., 2005). Therefore, we suggest that this excess silica is mainly derived from diagenetic chert (i.e., authigenic silica), and the sepiolite is a minor contribution of excess silica. The authigenic silica SiO2(auth) is additional evidence of upwelling (e.g., Beauchamp and Baud, 2002). The SiO2(auth) values in this study range from 0.00 wt.% to 44.88 wt.%, with an average of 27.72 wt.% (Table 1 and Fig. 1). It shows a rapid increasing to 40 wt.% from 0.00 wt.% in the lower part of stinkstones member in the Chihsia Formation and gradual decreasing to 25 wt.% in the upper part of the stinkstone member (Fig. 1). Oceanic upwelling carries the cold dissolved silica-charged water, hindering the silica dissolution and thus accounting for the supply of authigenic silica (Beauchamp and Baud, 2002). The Enshi Section was located near the paleowest of South China Block and thus affected by the eastern boundary current that was part of a circulation gyre within the Figure 4. The chert bands in the calcareous shale. The hammer as scale. Figure 5. The silty dolomite minerals in the dolostone intercalation. Sample ES28. The yellow bar 500 µm as scale. Paleotethys Ocean (Kutzbach and Guetter, 1990), which was usually associated with strong upwelling along the South China Block. This oceanic upwelling reasonably accounts for the rapid increasing of SiO2(auth) in the lower part of stinkstone member in the Chihsia Formation within the Early–Middle Gungurian. 4 CONCLUSIONS The stinkstones in the lower Chihsia Formation at Enshi Section in South China recorded a 3‰-magnitude negative excursion of bulk organic carbon isotope and a rapid increasing of authigenic silica content, suggesting a cause of oceanic upwelling in the western margin of PaleoTethys Ocean during the Early-Middle Kungurian stage. This upwelling also accounts for the widespread high organic carbon content sediments of the lower Chihsia Formation across the South China Block. ACKNOWLEDGMENTS We thank Jiaxin Yan and Wei Wang for their constructive comments. Research by H. Y. Wei is supported by the National Natural Science Foundation of China (No. 41302021), and by the Science and Technology Research Project of Jiangxi Province Education Department (No. GJJ13452). 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