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: hywei@ecit.edu.cn; weihengye@163.com
© 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). We also thank
Allison Young for improving this paper. Research by Hao Yu is
supported by the National Natural Science Foundation of China
(No. 41290260) and by the Ministry of Education of China (No.
20120001110052).
REFERENCES CITED
Algeo, T. J., Henderson, C. M., Tong, J., et al., 2013. Plankton
and Productivity during the Permian–Triassic Boundary
Crisis: An Analysis of Organic Carbon Fluxes. Global and
Planetary Change, 105: 52–67
Beauchamp, B., Baud, A., 2002. Growth and Demise of Permian Biogenic Chert along Northwest Pangea: Evidence
for End-Permian Collapse of Thermohaline Circulation.
Palaeogeography, Palaeoclimatology, Palaeoecology,
184: 37–63
Berner, R. A., 2002. Examination of Hypotheses for the
Permo-Triassic Boundary Extinction by Carbon Cycle
Modeling. Proceedings of National Academic Science
(USA), 99: 4172–4177
Bohacs, K., Suter, J., 1997. Sequence Stratigraphic Distribution
of Coaly Rocks: Fundamental Controls and Paralic Examples. American Association of Petroleum Geologists Bulletin, 81: 1612–1639
Cecil, C. B., 1990. Paleoclimatic Controls on Stratigraphic
Repetition of Chemical Siliciclastic Rocks. Geology, 18:
533–536
Des Marais, D. J., Strauss, H., Summons, R. E., et al., 1992.
Carbon Isotope Evidence for the Stepwise Oxidation of
the Proterozoic Environment. Nature, 359: 605–609
Feng, Z. Z., Yang, Y. Q., Jin, Z. K., 1997. Lithofacies and Palaeography of the Permian of South China. Petroleum
University Press, Beijing. 242 (in Chinese with Enghlish
Abstract)
Fenton, S., Grice, K., Twitchett, R. J., et al., 2007. Changes in
Biomarker Abundances and Sulfur Isotopes of Pyrite
A Kungurian Oceanic Upwelling on Yangtze Platform Evidenced by δ13Corg and Authigenic Silica
across the Permian-Triassic (P/Tr) Schuchert Dal Section
(East Greenland). Earth and Planetary Science Letters,
262: 230–239
Fielding, C. R., Frank, T. D., Birgenheier, L. P., et al., 2008.
Stratigraphic Imprint of the Late Paleozoic Ice Age in
Eastern Australia: A Record of Alternating Glacial and
Nonglacial Climate Regime. Journal of the Geological
Society, 165: 129–140
Frakes, L. A., Francis, J. E., Syktus, J. I., 1992. Climate Modes
of the Phanerozoic: The History of the Earth’s Climate
over the Past 600 Million Years. Cambridge University
Press, Cambridge. 274
Gastaldo, R. A., DiMichele, W. A., Pfefferkorn, H. W., 1996.
Out of the Icehouse into the Greenhouse: A Late Paleozoic
Analogue for Modern Global Vegetational Change. GSA
Today, 10: 1–7
Grard, A., François, L. M., Dessert, C., et al., 2005. Basaltic
Volcanism and Mass Extinction at the Permo-Triassic
Boundary: Environmental Impact and Modeling of the
Global Carbon Cycle. Earth and Planetary Science Letters,
234: 207–221
Hadjira, B., Wang, X., Ma, Z., et al., 2011. Preliminary Mineralogical and Geochemical Analysis on the Chihsia Formation of Tieqiao Section, Laibin, Guangxi and Their
Geological Implications. Geological Science and Technology Information, 30(1): 15–19 (in Chinese with English Abstract)
Hansen, H. J., 2006. Stable Isotopes of Carbon from Basaltic
Rocks and Their Possible Relation to Atmospheric Isotope
Excursions. Lithos, 92: 105–116
Hasiotics, S. T., Honey, J. G., 2000. Paleohydrologic and
Stratigraphic Significance of Crayfish Burrows in Continental Deposits: Examples from Several Paleocene Laramide Basin in the Rocky Mountains. Journal of Sedimentary Research, 70: 127–139
Hayes, J. M., Strauss, H., Kaufman, A. J., 1999. The Abundance of 13C in Marine Organic Matter and Isotopic Fractionation in the Global Biogeochemical Cycle of Carbon
during the Past 800 Ma. Chemical Geology, 161: 103–125
Heeremans, M., Larsen, B. T., Stel, H., 1996. Paleostress Reconstruction from Kinematic Indicators in the Oslo Graben, Southern Norway: New Constraints on the Mode of
Rifting. Tectonophysics, 266: 55–79
Hunt, J. M., 1996. Petroleum Geochemistry and Geology, 2nd
Edition. Freeman and Company, New York. 743
Jones, A. T., Frank, T. D., Fielding, C. R., 2006. Cold Climate
in the Eastern Australian Mid to Late Permian may Reflect
Cold Upwelling Waters. Palaeogeography, Palaeoclimatology, Palaeoecology, 237: 370–377
Kaufman, A. J., Knoll, A. H., Narbonne, G. M., 1997. Isotopes,
Ice Ages, and Terminal Proterozoic Earth History. Proceedings of the National Academy of Science (USA), 94:
6600–6660
Kiehl, J. T., Shields, C. A., 2005. Climate Simulation of the
Latest Permian: Implications for Mass Extinction. Geology, 33: 757–760
Korte, C., Kozur, H. W., 2010. Carbon-Isotope Stratigraphy
across the Permian-Triassic Boundary: A Review. Journal
217
of Asian Earth Sciences, 39: 215–235
Korte, C., Pande, P., Kalia, P., et al., 2010. Massive Volcanism
at the Permian-Triassic Boundary and Its Impact on the
Isotopic Composition of the Ocean and Atmosphere.
Journal of Asian Earth Sciences, 37: 293–311
Kutzbach, J. E., Guetter, P. J., 1990. Simulated Circulation of
an Idealized Ocean for Pangaean Time. Paleoceanography,
5: 299–317
Liu, X. Y., Yan, J. X., 2007. Nodular Chert of the Permian
Chihsia Formation from South China and Its Geological
Implications. Acta Sedimentologica Sinica, 25: 730–736
(in Chinese with English Abstract)
Lu, B. Q., Qu, J. Z., 1989. Anoxic Deposition Formed under
Upwelling and Transgression during the Early Permian of
South China. Chinese Science Bulletin, 35: 1193–1198
Ma, X. H., Zhang, Z. K., 1986. Palaeomagnetism and Its Use in
Search of Plate Tectonics. In: Li, C. Y., ed., On Principle
Problems of Plate Tectonics. Seismology Publishing
House, Beijing. 119–142 (in Chinese)
McLennan, S. M., 2001. Relationships between the Trace Element Composition of Sedimentary Rocks and Upper Continental Crust. Geochemistry, Geophysics, Geosystems, 2:
2000GC000109
Meyer, K. M., Yu, M., Lehrmann, D., et al., 2013. Constraints
on Early Triassic Carbon Cycle Dynamics from Paired
Organic and Inorganic Carbon Isotope Records. Earth and
Planetary Science Letters, 361: 429–435
Meyers, P. A., 1997. Organic Geochemical Proxies of Paleoceanographic, Paleolimnologic, and Paleoclimatic Processes. Organic Geochemistry, 27: 213–250
Mii, H. S., Shi, G. R., Cheng, C. J., et al., 2012. Permian
Gondwanaland Paleoenvironment Inferred from Carbon
and Oxygen Isotope Records of Brachiopod Fossils from
Sydney Basin, Southeast Australia. Chemical Geology,
291: 87–103
Montañez, I. P., Poulsen, C. J., 2013. The Late Paleozoic Ice
Age: An Evolving Paradigm. The Annual Review of Earth
and Planetary Science, 41: 629–656
Olaussen, S., Larsen, B. T., Steel, R., 1994. The Upper
Carboniferous–Permian Oslo Rifting: Basin Fil in Relation to Tectonic Development. In: Embry, A., ed., Pangea:
Global Environments and Resources. Canadian Society of
Petroleum Geology, 17: 175–197
Popp, B. N., Parekh, P., Tilbrook, T., et al., 1997. Organic Carbon δ13C Variations in Sedimentary Rocks as
Chemostratigraphic and Paleoenvironmental Tools. Palaeogeography, Palaeoclimatology, Palaeoecology, 132:
119–132
Powell, M. G., Schöne, B. R., Dorrit, E. J., 2009. Tropical Marine Climate during the Late Paleozoic Ice Age Using
Trace Element Analysis of Brachiopods. Palaeogeography,
Palaeoclimatology, Palaeoecology, 280: 143–149
Retallack, G. J., Jahren, A. H., 2008. Methane Release from
Igneous Intrusion of Coal during Late Permian Extinction
Events. The Journal of Geology, 116: 1–20
Schoepfer, S. D., Henderson, C. M., Garrison, G. H., et al.,
2013. Termination of a Continent-Margin Upwelling System at the Permian–Triassic Boundary (Opal Creek, Al-
218
berta, Canada). Global and Planetary Change, 105: 21–35
Schwab, V., Spangenberg, J. E., 2004. Organic Geochemistry
across the Permian-Triassic Transition at the Idrijca Valley,
Western Slovenia. Applied Geochemistry, 19: 55–72
Shen, J., Algeo, T., Hu, Q., et al., 2013. Volcanism in South
China during the Late Permian and Its Relationship to
Marine Ecosystem and Environmental Changes. Global
and Planetary Change, 105: 121–134
Shen, S. Z., Wang, Y., Henderson, C. M., et al., 2007. Biostratigraphy and Lithofacies of the Permian System in the
Laibin-Heshan Area of Guangxi, South China. Palaeoworld, 16: 120–139
Shi, G. R., 1995. The Late Palaeozoic Brachiopod Genus
Yakovlevia Fredericks, 1925 and the Yakovlevia Transversa Zone, Northern Yukon Territory, Canada. Proceedings of the Royal Society of Victoria, 107: 51–71
Shi, G. R., Grunt, T. A., 2000. Permian Gondwana-Boreal Antitropicality with Special Reference to Brachiopod Faunas.
Palaeogeography, Palaeoclimatology, Palaeoecology,
155: 239–263
Svensen, H., Planke, S., Malthe-Sørenssen, A., et al., 2004.
Release of Methane from a Volcanic Basin as a Mechanism for Initial Eocene Global Warming. Nature, 429:
542–545
Svensen, H., Planke, S., Polozov, A. G., et al., 2009. Siberian
Gas Venting and the End-Permian Environmental Crisis.
Earth and Planetary Science Letters, 277: 490–500
Tong, J. N., Shi, G. R., 2000. Evolution of the Permian and
Triassic Foraminifera in South China. In: Yin, H. F., Dickins, J. M., Shi, G. R., et al., eds., Permian–Triassic Evolution of Tethys and Western Circum-Pacific. Elservier, Amsterdam. 291–307
Trappe, J., 1994. Pangean Phosphorites-Ordinary Phosphorite
Genesis in an Extraordinary World? Canadian Society of
Petroleum Geologists, Memoir, 17: 469–478
Walker, B. D., Guilderson, T. P., Okimura, K. M., et al., 2014.
Radiocarbon Signatures and Size-Age-Composition Rela-
Hao Yu and Hengye Wei
tionships of Major Organic Matter Pools within a Unique
California
Upwelling
System.
Geochimica
et
Cosmochimica Acta, 126: 1–17
Wang, Y., Jin, Y. G., 1998. Permian Topographic Evolution of
the Jiangnan Basin, South China. In: Retanasthien, B.,
Rieb, S. L., eds., Proceedings of the International Symposium on Shallow Tethys 5. Chiang Mai University Press,
Chiang Mai, Thailand. 1–497
Wei, H., Chen, D., 2011. Lithofacies Palaeogeography of the
Qixia Age of Permian in Western Hubei-Northwestern
Hunan Provinces. Journal of Palaeogeography, 13:
551–562 (in Chinese with English Abstract)
Wei, H., Chen, D., Wang, J., et al., 2012. Organic Accumulation
in the Lower Chihsia Formation (Middle Permian) of
South China: Constraints from Pyrite Morphology and
Multiple Geochemical Proxies. Palaeogeography, Palaeoclimatology, Palaeoecology, 353–355: 73–86
Werne, J. P., Hollander, D. J., 2004. Balancing Supply and Demand: Controls on Carbon Isotope Fractionation in the
Cariaco Basin (Venezuela) Younger Dryas to Present. Marine Chemistry, 92: 275–293
Winguth, A. M. E., Heinze, C., Kutzbach, J. E., et al., 2002.
Simulated Warm Polar Currents during the Middle Permian.
Paleoceanography,
17(4):
911–918
doi:10.1029/2001PA000646
Whiticar, M. J., 1996. Stable Isotope Geochemistry of Coals,
Humic Kerogens and Related Natural Gases. International
Journal of Coal Geology, 32: 191–215
Yan, J. X., Munnecke, A., Steuber, T., et al., 2005. Marine Sepiolite in Middle Permian Carbonates of South China: Implications for Secular Variation of Phanerozoic Seawater
Chemistry. Journal of Sedimentary Research, 75: 328–338
Yin, H. F., Jiang, H. S., Xia, W. C., et al., 2014. The
End-Permian Regression in South China and Its Implication on Mass Extinction. Earth-Science Reviews, 137:
19–33