A preview of this full-text is provided by Cambridge University Press.
Content available from Geological Magazine
This content is subject to copyright.
Geol. Mag. 155 (8), 2018, pp. 1667–1683. c
Cambridge University Press 2017 1667
doi:10.1017/S0016756817000462
Carbon isotope perturbations and faunal changeovers during the
Guadalupian mass extinction in the middle
Yangtze Platform, South China
HENGYE WEI∗†‡, QUZONG BAIMA†, ZHEN QIU§‡& CHAOCHENG DAI∗
∗State Key Laboratory Breeding Base of Nuclear Resources and Environment, East China University of Technology,
Nanchang, Jiangxi Province, 330013, China
†School of Earth Science, East China University of Technology, Nanchang, Jiangxi Province, 330013, China
§PetroChina Research Institute of Petroleum Exploration and Development, Beijing 100083, China
(Received 17 January 2017; accepted 3 May 2017; first published online 5 June 2017)
Abstract – The Guadalupian mass extinction took place during the major global environmental
changes during Phanerozoic time. Large-scale sea-level fluctuations and a negative shift of δ13Cwere
associated with this crisis. However, the diagenetic or primary origin of the decreased δ13C across the
Guadalupian–Lopingian (G–L) boundary and the potential causes for this biotic crisis are still being
intensely debated. Integrated analyses, including detailed petrographic examination, identification of
foraminifer and fusulinid genera, and analysis of carbonate δ13Ccarb and bulk δ13 Corg across the G–L
boundary were therefore carried out at Tianfengping, Hubei Province, South China. Our results show
that: (1) some foraminifer and most fusulinid genera disappear in the upper Maokou Formation (upper
Guadalupian); (2) the negative shift of δ13Ccarb in the uppermost Maokou Formation is of diagenetic
origin, but the values of δ13Ccarb in the remainder of the Maokou Formation and in the Wuchiap-
ing Formation represent a primary signal of coeval seawater; and (3) the bulk δ13Corg perturbation
across the G–L boundary at Tianfengping is mainly controlled by organic matter (OM) source, that is,
terrestrial OM contribution. We suggest that the δ13Ccarb negative shift in the lower Wuchiaping Form-
ation (Wuchiapingian) compared to that in the lower–middle Maokou Formation (Capitanian) were
probably caused by the re-oxidization of 12C-rich OM during regression. Global regression resulted in
the negative shift of δ13Ccarb at the G–L boundary in South China and led to the loss of shallow-marine
benthic habitat. Large-scale global regression is probably one of the main causes for this bio-crisis.
Keywords: Guadalupian–Lopingian, foraminifer, carbon isotope, mass extinction, South China.
1. Introduction
A mass extinction occurred in the Guadalupian epoch
(Jin, Zhang & Shang, 1994; Stanley & Yang, 1994;
Clapham, Shen & Bottjer, 2009), called the end-
Guadalupian mass extinction or pre-Lopingian crisis
(Jin, Zhang & Shang, 1994; Stanley & Yang, 1994;
Shen & Shi, 1996,2002; Wang & Sugiyama, 2000)or
mid-Capitanian mass extinction (Wignall et al. 2009b;
Bond et al. 2010,2015). This bio-crisis affected mar-
ine taxa including fusulinids, small foraminifers, cor-
als, brachiopods, bivalves and ammonoids (Jin, Zhang
& Shang, 1994; Wang & Sugiyama, 2000; Weidlich,
2002; Isozaki & Aljinovi´
c, 2009;Weiet al. 2012;
Hada et al. 2015; Zhang, Wang & Zheng, 2015).
Several geological events have been proposed as the
main cause of the mass extinction, including Emeishan
volcanism (Zhou et al. 2002; Wignall et al. 2009a;
Sun et al. 2010), large-scale sea-level fall and loss
of shallow-marine habitat (Chen, George & Yang,
2009; Wignall et al. 2009b;Qiuet al. 2014), cool-
‡Authors for correspondence: weihengye@163.com; qiuzhen316@
163.com
ing (Isozaki, Kawahata & Minoshima, 2007; Isozaki,
Aljinovic & Kawahata, 2011; Kofukuda, Isozaki &
Igo, 2014) and marine anoxia (Isozaki, 1997; Saitoh
et al. 2013b; Zhang et al. 2015;Weiet al. 2016). How-
ever, the causes for this biocrisis are still disputed.
There is a negative excursion of carbon isotope as-
sociated with this mass extinction at the Guadalupian–
Lopingian (G–L, 259.1 ±0.5 Ma, Zhong et al. 2014)
boundary (Wang, Cao & Wang, 2004; Wignall et al.
2009a; Bond et al. 2010). Wignall et al. (2009a) inter-
preted the negative excursion of carbon isotope at the
G–L boundary as the carbon cycle perturbation resul-
ted from Emeishan volcanism. Further, they sugges-
ted that the negative excursion discovered by Wang,
Cao & Wang (2004) actually occurred during middle
Capitanian time (see also Bond et al. 2010). How-
ever, Nishikane et al. (2014) questioned the volcanic
mechanism and middle Capitanian negative excursion
of carbon isotope. They argued that the drop in eu-
static sea level during Guadalupian time would not
be consistent with widespread volcanism since en-
hanced volcanism is generally associated with a high
plate production rate which would result in sea-level
rise. The high carbon isotope ratios during middle
1668 H. WEI AND OTHERS
Capitanian time at Penglaitan global boundary stra-
totype section and point (GSSP) section (see Chen
et al. 2011) is also inconsistent with the negative excur-
sion of carbon isotope during middle Capitanian time
(Nishikane et al. 2014). Instead, they suggested a de-
clined primary production as the cause for the negat-
ive excursion at the G–L boundary (Nishikane et al.
2014; see also Yan, Zhang & Qiu, 2013). A new in-
terpretation for the negative excursion of carbon iso-
tope at the G–L boundary was suggested by Saitoh
et al. (2013a) who interpreted it as the result of the up-
welling of oxygen-depleted water rich in 12C onto the
euphotic shelf. Alternatively, Jost et al. (2014) ques-
tioned the reliability of carbon-isotope negative excur-
sion at the G–L boundary as a primary signal, and in-
terpreted it as local burial conditions or diagenetic ori-
gin in some important sections. Accordingly, the inter-
pretation for the carbon isotope negative excursion at
the G–L boundary has been highly controversial, and
needs more work to reveal the causes of the carbon iso-
tope changes and the kill-mechanism of this extinction.
Combining detailed petrographic analysis via thin-
section, we have analysed facies, foraminifer fossils
record, carbonate-carbon and bulk organic-carbon iso-
tope changes across the G–L boundary at the Tian-
fengping section in Enshi city (in Hubei Province) in
the middle Yangtze Platform, South China. Our results
show a different interpretation for the carbon isotope
changes during the boundary interval.
2. Geological background
The road-side Tianfengping section (30°1937 N,
109°1852 E) is located at the Tianfengping village
in Enshi city in western Hubei Province, South China.
The Tianfengping section crops out over the Maokou
Formation, Kuhfeng Formation and Wuchiaping Form-
ation, in ascending order. A detailed description of li-
thology and interpretation is provided in Section 4.
Located in the eastern Palaeo-Tethys ocean in the trop-
ical zone (Scotese & Langford, 1995, p. 3; Muttoni
et al. 2009), the South China Block was during Cap-
itanian time a large carbonate platform divided by
the deep-water Jiangnan Basin in the middle into the
Yangtze Platform in the west and the Cathaysian Plat-
form in the east (Fig. 1). The Xiakou-Lichuan Bay
(Yin et al. 2014) was located in the northern Yangtze
Platform (Fig. 1). Our studied section at Tianfengping
was located in the centre of this bay. The Kangdian old
land was located in the west of the Yangtze Platform.
From middle Capitanian time, the Emeishan large
igneous province (LIP) erupted in the southwestern
Yangtze Platform (Ali et al. 2005; Wignall et al.
2009a), resulting in a large volcanic and volcan-
iclastic succession which accumulated across the
G–L boundary. The volume of Emeishan LIP was
from 0.3×106km3(Xu et al. 2001)to0.6×106km3
(Yin et al. 1992, p. 146), only about one-tenth
of the Siberian LIP with a volume of 4×106km3
(Courtillot, 1999). Several rift basins including the Qi-
anzhong basin near Guiyang in Figure 1 (Chen et al.
2003) and the Xiakou-Lichuan Bay were suggested to
be of rift origin by Zhu (1989), which may be related
to the thermal decay of the Emeishan plume.
3. Methods
One hundred samples were collected at Tianfengping
with a c. 25 cm sample interval, avoiding weathered
samples. One hundred thin-sections were created for
petrographic examination and identification of fossils.
For inorganic-carbon isotope measurements, we pre-
pared 57 samples of bulk carbonate rock. For each
sample, a fresh chip was ground using an agate mor-
tar. Powdered samples were dissolved in phosphoric
acid to release CO2at Kiel IV of automated carbon
reaction device, which was coupled with a Finnigan
MAT 253 mass spectrometer for δ13C and δ18 O meas-
urements. All C-isotope ratios are calibrated to V-
PDB using NBS-19. Analytical precision for δ13C and
δ18Ois±0.04 ‰and ±0.08 ‰(1σ), respectively. This
experiment was carried out at the Nanjing Institute
of Geology and Palaeontology, Chinese Academy of
Sciences.
For bulk organic-carbon isotope analyses, 84
samples were powdered smaller than 200 mesh using
an agate ball mortar. These powdered samples were
digested by 6 N HCl to remove all carbonates. The
acid-insoluble residues were cleaned and dried, mixed
with CuO and sealed in vacuo for further furnace
processing. The samples were combusted at 800 °C
and the released CO2was cryogenically extracted and
sealed in vacuum tubes for subsequent 13C/12C determ-
ination using a Finnigan MAT 253 at the Nanjing Insti-
tute of Geology and Palaeontology, Chinese Academy
of Sciences. Reproducibility was better than ±0.08 ‰
for organic carbon calibrated to a urea (IVA33802174)
standard with δ13Corg value of –40.73 ‰. All data are
reported in per mille (‰) relative to V-PDB standard.
4. Results
4.a. Lithostratigraphy
The Tianfengping section consists of the middle
Permian Maokou and Kuhfeng formations, and
the upper Permian Wuchiaping Formation (Fig. 2).
The Wuchiaping Formation can be subdivided
into the Wangpo Shale Member in the lower part and
the Xiayao Limestone Member in the upper part (Feng,
Yang & Jin, 1997; p. 75).
4.a.1. Maokou Formation
The Maokou Formation is composed of massive lime-
stones/dolostones and is subdivided into light-grey
limestones in the lower part, grey to dark-grey lime-
stones in the middle part and limy dolostones in the
upper part (Figs 2,3a). The limestones in the lower and
middle Maokou Formation show abundant stylolites
C-isotope and faunal changeover across GLB 1669
Figure 1. (Colour online) Capitanian palaeogeography of South China (modified from Zhu, 1989;Wang&Jin,2000;Chenet al.2003;
Du et al.2015) and the locations of studied sections. TP – Tianfengping section; QZ Basin – Qianzhong intrashelf basin; JN Basin –
Jiangnan basin; TP – Tianfengping section; TQ – Tieqiao section.
and bioturbation (Fig. 2), and contain abundant bio-
clasts such as brachiopods, crinoids, echinoids, fo-
raminifers, sponge spicules and green calcareous algae
(Figs 2,3b). Non-skeletal grains such as peloids, onc-
oids and cortoids also occur in the lower and middle
Maokou Formation. The limy dolostones in the up-
permost Maokou Formation contain dolomite crystals
with cloudy centres and clear rims (Fig. 3c), suggest-
ing replacement of limestones. There are also abund-
ant bitumen (Fig. 3d) and small karst caves (Fig. 3e,
f) in the uppermost Maokou Formation, suggesting a
phase of karstification during sub-aerial exposure. The
boundary between the Maokou and Kuhfeng forma-
tions is a regional unconformity (Fig. 3a).
4.a.2. Kuhfeng Formation
The 3.8 m-thick Kuhfeng Formation is composed
of thin-bed (3–8 cm thick) black cherts interbedded
with c. 1 cm carbonaceous black shales (Fig. 3a).
Authigenic carbonate concretions are common and
the largest (c. 1 m) occurs in the middle Kuh-
feng Formation (Fig. 3g). Microscopically, phylloid al-
gae (Fig. 3h), siliceous sponge spicules (Fig. 3i) and
abundant radiolarians (Fig. 3j) occur in the lower,
middle and upper parts of Kuhfeng Formation, re-
spectively. Some small ammonoids (Fig. 3k) and bra-
chiopods (Fig. 3l) also occur in the middle Kuhfeng
Formation.
4.a.3. Wuchiaping Formation
The Wuchiaping Formation includes the Wangpo
Shale Member in the lower part and Xiayao Limestone
Member in the upper part. The Wangpo Shale Member
consists of lithic arenite and claystones in the lower
part and black shale intercalated with thin-bed lime-
stone in the upper part (Fig. 2). Coal is very common in
the lower Wangpo Shale Member (Fig. 4a). The lithic
arenite contains abundant feldspar and pyrite miner-
als (Fig. 4b–d) and lithic grains including tuff (Fig. 4e,
f), basalt (Fig. 4g), spherulitic rhyolite (Fig. 4h) and
chert (Fig. 4i). The tuff grains are very common. Mus-
covite (Fig. 4j), authigenic gypsum (Fig. 4k) and albite
(Fig. 4l) also occur in this sandstone. The claystones
overlying the lithic arenite are rich in pyrites (Fig. 4m),
and the overlying floatstone contains abundant phyl-
loid green algae (Fig. 4n). The thin-bedded lime mud-
stones occurred as intercalated bed in the black shale in
the upper Wangpo Shale Member, containing abundant
small round peloids (Fig. 4o).
The Xiayao Limestone Member in the upper
Wuchiaping Formation consists of thin- to thick-
bedded argillaceous lime mudstones, wackestones and
packstones (Fig. 5a, b). The argillaceous lime mud-
stones and wackestones display laminations (Fig. 5c)
and contain large brachiopods (Fig. 5d–f), partially re-
placed by authigenic albite (Fig. 5e). The packstones
to grainstones in the upper Xiayao Limestone Mem-
ber contain abundant green algae and sponge spicules,
gastropods, brachiopods, foraminifers (Fig. 5g–i) and
ubiquitous disseminated glacuconites (Fig. 5h).
4.b. Foraminifer biostratigraphy
Fossil range data show that the Maokou Formation
contains a high diversity of nonfusulina foraminifers
1670 H. WEI AND OTHERS
Figure 2. (Colour online) Graphic sedimentary log across the
Guadalupian–Lopingian boundary at the Tianfengping section,
South China. S – shale; M – lime mudstone; W – wackestone; P
– packstone; G – grainstone; Sa – sandstone. The conodont data
is from Xia et al.(2006). C. p.p. –Clarkina postbitteri postbit-
teri;C.d. –Clarkina dukouensis;C.a. –Clarkina asymmetrica.
and fusulinids, but that the Wuchiaping Formation
only contains a few small foraminifers at Tianfeng-
ping (Fig. 6). For the small foraminifers, the long-
ranging genera persisting into the Wuchiapingian
strata include Pachyphloia,Nodosaria,Langella,
Geinitzina,Psedoglandulina,Neotuberitina,Hemigor-
dius and Agathammina at Tianfengping (Fig. 6), in
which Pachyphloia,Nodosaria,Geinitzina,Hemig-
ordius and Agathammina had been also reported in
Wuchiapingian deposits in Guangyuan in South China
(Lai et al. 2008), at Tieqiao in South China (Wignall
et al. 2009b; Zhang et al. 2015) and in Takachiho in
Japan (Kobayashi, 2012). The Globivalvulina,Deck-
erella,Palaeotextularia,Cribrogenerina,Climacam-
mina,Tetrataxis,Frondicularia,Glomospira,Cribro-
generina,Neodiscus,Archaediscus,Multidiscus,Am-
modiscus,Plectogyra and Robuloides disappeared in
the Wuchiaping Formation at Tianfengping (Fig. 6).
However, some of these disappeared genera had been
reported in the Wuchiapingian strata elsewhere, such
as Palaeotextularia in Guangyuan in South China (Lai
et al. 2008), Climacammina in Guangyuan (Lai et al.
2008), Laibin (Wignall et al. 2009a; Zhang et al. 2015)
in South China and Takachiho in Japan (Kobayashi,
2012), Frondicularia in Guangyuan (Lai et al. 2008),
Laibin Wignall et al. (2009b)and Takachiho in Ja-
pan (Kobayashi, 2012), Glomospira in Takachiho in
Japan (Kobayashi, 2012), Neodiscus in Takachiho in
Japan (Kobayashi, 2012), and Multidiscus in Laibin
in South China (Zhang et al. 2015) and Takachiho
in Japan (Kobayashi, 2012). Most of the foraminiferal
genera therefore persist into the Wuchiapingian strati-
graphy and only nine of them disappear from the upper
Maokou Formation.
For the fusulinids range data, all of these fusulinid
genera identified in the Maokou Formation disap-
peared in the Kuhfeng and the Wuchiaping forma-
tions (Wuchiapingian) at Tianfengping (Fig. 6). They
were Schwagerina,Schubertella,Reichellina,Skin-
nerella,Ozawainella,Chenella,Chusenella,Staffella
and Nankinella.However,Reichellina,Codonofusi-
ella,Staffella and Nankinella had been reported in
Wuchiapingian deposits elsewhere such as Guangy-
uan (Lai et al. 2008), Laibin Wignall et al. (2009b)
in South China (Jin et al. 2006) and in Takachiho in
Japan (Kobayashi, 2012). Most of the fusulinid gen-
era therefore disappear at 9.5 m in the upper Maokou
Formation in the upper Capitanian (e.g. J. granti zone;
Xia et al. 2006) at Tianfengping (Fig. 6).
4.c. Carbon isotope chemostratigraphy
At Tianfengping, carbonate-carbon isotopic ratios
δ13Ccarb range from –0.7 ‰to 3.9 ‰with an average
value of 2.3 ‰(Tabl e 1 ). The δ13 Ccarb profile stabilizes
at c. 3.8 ‰in the lower and middle Maokou Formation
below 9.0 m and shifts to 0.7‰in the upper Maokou
Formation (Fig. 7). The δ13Ccarb values in the Wuchiap-
ing Formation are relatively low, ranging from 0.2 ‰
to 1.6 ‰with an average value of 0.8‰(Fig. 7). The
δ13Ccarb profile in the Wuchiaping Formation shows
a gradual positive change from c. 0.5 ‰to 1.60 ‰
(Fig. 7). Carbonate δ18Ocarb values range from –7.0 ‰
to –3.5 ‰, with an average value of –5.6 ‰(Ta b l e 1).
Organic-carbon isotopic ratios (δ13Corg) range from
–28.7 ‰to –21.5 ‰, with an average value of –
26.0 ‰. The δ13Corg profile stabilizes at c. –28.4 ‰
in the lower and middle Maokou Formation (below
9.0 m) and shows a positive change from –28.4 ‰to
c. –26.5 ‰in the upper Maokou Formation (Fig. 7).
The δ13Corg profile stabilizes at c. –26.5 ‰in the
lower Kuhfeng Formation and changes to c. –27.1 ‰
in the upper Kuhfeng Formation. At the Kuhfeng–
Wuchiaping formation boundary, δ13Corg profile shows
an abrupt shift from –26.5 ‰in the Kuhfeng Form-
ation to –21.5 ‰in the lower Wangpo Shale Mem-
ber. The δ13Corg profile then stabilizes at c. –23.9 ‰
in the upper Wangpo Shale Member and lower Xiayao
C-isotope and faunal changeover across GLB 1671
Figure 3. (Colour online) Field photos and thin-section micrographs of the Maokou and Kuhfeng formations at Tianfengping, South
China. (a) The Maokou–Kuhfeng boundary; and (b) Udoteacean packstone. Sample TP18. Plane polarized light (PPL). Bar scale
500 μm. (c) Dolomites with cloudy centres and clear rims. Sample TP42. PPL. Bar scale 500 μm. (d) Bitumens (arrows) at the top
of the Maokou Formation. Size of bitumen is about 1 cm. (e) Small karst caves (dash lines) at the top of the Maokou Formation. (f)
The internal surface of karst cave (dash line) the top of the Maokou Formation. Pencil for scale. (g) Dolomite nodule in the black
chert. Hammer for scale. (h) Phylloids in the black chert in the lowermost Kuhfeng Formation. Sample TP45. PPL. (i) Siliceous
sponge spicule (arrow) in the black chert. Sample TP50. PPL. (j) Radiolarian chert. Sample TP55. PPL. (h–j) Bar scale 500 μm. (k)
Ammonoid in the lower Kuhfeng Formation. Sample TP46. PPL. Bar scale 50 μm. (l) Small and thin-shell brachiopod fragments
(solid arrows) and intact one (hollow arrow). Sample TP50. PPL. Bar scale 200 μm.
Limestone Member, and then gradually shifts to –
25.8 ‰in the upper Xiayao Limestone Member
(Fig. 7).
In summary, δ13Ccarb values in the Wuchiaping
Formation are substantially lower than in the Maokou
Formation (c. 3‰in magnitude). The δ13Ccarb val-
ues in the upper Maokou Formation are isotopically
lighter than in the lower and middle Maokou Forma-
tion. The δ13Corg values in the upper Maokou Form-
ation and Kuhfeng Formation are higher than in the
lower and middle Maokou Formation, while δ13Corg
values in the Wuchiaping Formation are much higher
than in the Kuhfeng and Maokou formations. Even in
the Wuchiaping Formation, the δ13Corg profile displays
1672 H. WEI AND OTHERS
Figure 4. (Colour online) Field photos and thin-section micrographs of the Wangpo Shale Member in the Wuchiaping Formation
at Tianfengping. (a) Black coal (arrow). Marker for scale. (b) Feldspar (F), tuff fragment (TF) and pyrite (P) in sandstone. Sample
TP59. PPL. Bar scale 500 μm. (c) Cross-polarized light (CPL) of (b). (d) Palgioclase in the sandstones. Bar sample TP 59. CPL.
(e) Tuff fragments (TF) showing oxidized rim. Sample TP59. PPL. (f) Tuff fragment (TF) showing glassy fragment (arrow). Sample
TP59. PPL. (g) Basaltic fragment (BF) showing angular shape. Sample TP60. CPL. (h) Spherulitic rhyolite fragment (RF). Sample
TP59. PPL. (i) Chert fragment (CF) and feldspar (F) grains. Sample TP60. CPL. (j) Muscovite (arrow) grain. Sample TP60. CPL.
(k) Authigenic gypsum. Sample TP59. CPL. (l) Authigenic albite in the sandstone. Sample TP62. CPL. (d–l) Bar scale 100 μm. (m)
Claystones containing abundant black pyrite. Sample TP67. PPL. (n) Phylloid fragments (light bands) and calcisphere (light spots).
Sample TP69. PPL. (o) Peloid wackestone. Sample TP74. PPL. (m–o) Bar scale 500 μm.
C-isotope and faunal changeover across GLB 1673
Figure 5. (Colour online) Field photos and thin-section micrographs of the Xiayao Limestone Member in the Wuchiaping Formation at
Tianfengping. (a) Outcrop of the Wuchiaping Formation; (b) Lime mudstone, Sample TP 84. PPL. (c) Laminations in the argillaceous
lime mudstone. Sample TP83. PPL. (d) Large brachiopods (Br) in the argillaceous lime mudstone. Sample TP86. PPL. (b–d) Bar scale
500 μm. (e) CPL image of (d). Authigenic albites occur in the rim of brachiopod fragments. (f) Brachiopod fossil in Sample TP93 in
Bed 17. (g) Gymnocodiacean algae (Gy) grainstone containing small foraminifera (Fo), brachiopod spicule (Br) and gastropod (Ga).
Sample TP95. PPL. Bar scale 500 μm. (h) Disseminated glauconites (yellow-green color). Sample TP95. PPL. Bar scale 100 μm. (i)
Packstone containing abundant sponge spicules (Sp and hollow arrows), trilobite (Tr) and Gymnocodiacean green algae (Gy). Sample
TP97. PPL. Bar scale 500 μm.
three steps from heavier to lighter values in the lower
Wangpo Shale Member, upper Wangpo Shale Member
to lower Xiayao Limestone Member and upper Xiayao
Limestone Member, respectively.
5. Discussion
5.a. Changes in the sedimentary environment
The unconformity at the Maokou–Kuhfeng forma-
tion boundary at Tianfengping also occurred in other
sections, for example Jianshi and Badong, suggest-
ing a regional regression and sub-aerial exposure in
South China (e.g. Chen et al. 2000;Niuet al. 2000).
The occurrence of phylloids and small brachiopods
in the black chert of the Kuhfeng Formation sug-
gests that it was not a typical deep-marine environ-
ment. The spherical radiolarians which prefer inhab-
iting relatively shallower-water environments (Kozur,
1993) are abundant in the Kuhfeng Formation at Mao-
caojie near our studied section (Shi et al. 2016).
Furthermore, high organic carbon (6 %, Yao et al.
2002; 1.5–18 %, Shi et al. 2016) in the Capitanian
1676 H. WEI AND OTHERS
Table 1. Stable carbonate-carbon, bulk organic-carbon, oxygen
isotopes and the isotopic difference between carbonate-carbon and
organic-carbon isotope data at Tianfengping, South China.
Sample Thickness (m) δ13Corg (‰)δ13 Ccarb (‰)δ18Ocarb (‰)
Xiayao Limestone Member (Wuchiaping Formation)
TP100 31.95 −25.8 1.6 −5.4
TP99 31.4 −26.1 1.3 −5.5
TP98 31.05 −25.6 1.3 −5.6
TP97 30.75 −25.5 0.2 −5.3
TP96 29.75 −25.2 0.9 −5.9
TP95 29.55 −25.4 1.5 −4.5
TP94 28.75 −24.4 0.4 −6.8
TP93 28.45 −23.9
TP92 28.05 −24.1
TP91 27.55 −23.7
TP90 27.3 −23.8
TP89 26.75 −24.5 0.9 −4.5
TP88 26.45 −24.2 1.6 −5.6
TP87 26.15 −24.3 0.6 −4.5
TP86 25.75 −24.0
TP85 25.4 −25.6 0.5 −5.0
TP84 24.8 −24.4 0.7 −4.5
TP83 24.35 −23.9 0.4 −3.6
TP82 23.95 −23.9 0.2 −3.5
TP81 23.4 −23.4 0.3 −3.9
Wangpo Shale Member (Wuchiaping Formation)
TP80 22.85 −24.1
TP79 22.5 −24.5
TP78 22.1 −23.7
TP77 21.6 −24.4
TP76 21.1 −24.6
TP75 20.7 −24.2
TP74 20.3
TP73 20.1
TP72 19.6
TP71 19.3 −24.7
TP70 19
TP69 18.8 −24.7
TP68 18.5
TP67 18.2
TP66 17.95
TP65 17.65 −21.5
TP64 17.35 −22.8
TP63 17.15 −23.0
TP62 16.85 −23.9
TP61 16.55 −24.5
TP60 16.25 −23.1
TP59 15.95 −24.6
TP58 15.65 −23.7
Kuhfeng Formation
TP57 15.55 −26.7
TP56 15.35 −26.9
TP55 15.05 −27.2
TP54 14.75 −27.0
TP53 14.45 −26.9
TP52 14.15 −26.9
TP51 13.85 −27.1
TP50 13.7 −27.0
TP49 13.4 −26.7
TP48 13.1 −26.7
TP47 12.8 −27.1
TP46 12.4 −26.5
TP45 12.1 −26.7
Maokou Formation
TP44 11.75 −28.3 0.9 −6.7
TP43 11.5 −26.9 1.1 −6.4
TP42 11.2 −26.5 0.7 −7.3
TP41 10.9 −26.9 1.1 −6.4
TP40 10.6 −27.0 0.4 −7.0
TP39 10.3 −27.7 1.9 −5.4
TP38 10 −28.1 2.9 −6.0
TP37 9.8 −27.6 2.6 −6.1
TP36 9.5 −28.7 0.6 −4.7
TP35 9.2 −27.8 2.9 −5.8
TP34 8.9 −27.5 2.7 −6.5
Table 1. Continued.
Sample Thickness (m) δ13 Corg (‰)δ13Ccarb (‰)δ18 Ocarb (‰)
TP32 8.6 −28.5 2.8 −5.8
TP31 8.3 −28.3 3.2 −5.8
TP30 8 −27.3 2.4 −5.8
TP29 7.7 −28.4 3.3 −5.8
TP28 7.4 −28.5 3.3 −5.7
TP27 7.1 −28.4 3.3 −5.9
TP26 6.8 −28.7 2.8 −5.9
TP25 6.5 −28.2 2.8 −6.8
TP24 6.2 −28.4 3.5 −5.6
TP23 5.9 −28.5 3.2 −5.6
TP22 5.6 −28.2 3.4 −5.7
TP21 5.3 −27.8 3.0 −5.6
TP20 5.2 −28.3 3.5 −5.8
TP19 4.9 −28.2 3.2 −6.2
TP18 4.6 −28.2 3.4 −5.7
TP17 4.3 3.4 −5.7
TP16 4 −28.4 3.5 −6.0
TP15 3.7 3.6 −5.8
TP14 3.4 3.6 −5.7
TP13 3.1 3.4 −5.8
TP12 3 3.4 −5.6
TP11 2.8
TP10 2.5 3.6 −5.7
TP09 2.2
TP08 2.1 3.7 −5.7
TP07 1.8
TP06 1.5 3.6 −5.7
TP05 1.2 3.2 −5.7
TP04 0.9 3.5 −5.7
TP03 0.6 3.8 −5.6
TP02 0.3 3.9 −5.8
TP01 0 3.2 −5.9
Kuhfeng Formation at Tianfengping probably sug-
gests a restricted environment with weak water cir-
culation (e.g. Zhang et al. 2015;Weiet al. 2016;
Saitoh et al. 2017). The petrography, which is char-
acterized by abundant peloids, sandstones and clay-
stones intercalated with coal beds in the Wangpo
Shale Member (Wuchiapingian) (Fig. 4), suggests a
coastal swamp environment; the argillaceous lime-
stone and low biodiversity dominated by brachiopods
in the low Xiayao Limestone Member (Fig. 5b–f) sug-
gest a restricted environment, however. Abundant au-
thigenic gypsums and albites in the Wangpo Shale
Member and lower Xiayao Limestone Member sug-
gest high concentrations of SO42−and/or Na+in the
diagenetic fluids since the gypsums grew in inter-
granular pore (Fig. 4k) and the albites were formed
by the replacement of skeletons such as brachiopods
(Fig. 5e). In the upper Xiayao Limestone Member,
the packstones/grainstones contain abundant bioclasts
with high biodiverisity (Fig. 5g–i) and glauconites,
suggesting an open shallow-marine environment. In
summary, the lower part of the Wuchiaping Forma-
tion (upper Permian) was deposited in a coastal swamp
environment. The middle Permian Maokou Formation
and the upper part of Wuchiaping Formation (upper
Permian) were deposited in shallow-marine environ-
ments. The middle Permian Kuhfeng Formation, which
is sandwiched between the Maokou and Wuchiaping
formations, may have been deposited in a mid-depth
C-isotope and faunal changeover across GLB 1677
Figure 8. (Colour online) (a) Cross-plot between carbonate-carbon isotope and oxygen isotope at Tianfengping; (b) Cross-plot between
carbonate-carbon isotope and organic-carbon isotope.
environment rather than a deep basin. The Kuhfeng
Formation elsewhere in South China was also depos-
ited in the water depth not deeper than several hundred
metres (Kametaka et al. 2005).
5.b. Primary or diagenetic origin of carbon isotope records
5.b.1. δ13Ccarb changes
Carbon isotopic ratios can be altered by the post-
depositional processes such as meteoric burial and
organic diagenesis (e.g. Rosales, Quesada & Robles,
2001). It is critical to assess the diagenetic effect
on carbon isotopic composition in order to recon-
struct the environmental changes. The oxygen iso-
tope ratios of whole-rock carbonates are generally
altered during diagenesis (e.g. Weissert, Joachimski
& Sarnthein, 2008). Positive correlations between
δ18O and δ13 C in marine carbonate sediments are
often taken as evidence for diagenetic alteration,
as it is difficult to produce such arrays in a
primary depositional environment (Marshall, 1992;
Melim, Swart & Eberli, 2004; Knauth & Kennedy,
2009; Preto, Spotl & Guaiumi, 2009). At Tianfeng-
ping, there is a weak covariation between δ13Ccarb
and δ18Ocarb values (R2=0.25, Fig. 8a) in the
Maokou Formation, recording a primary seawater
signal. However, the negative shift of δ13Ccarb pro-
file at the top of the Maokou Formation is associ-
ated with the similar negative shift of δ18Ocarb pro-
file (Fig. 7), suggesting a diagenetic origin in this
interval although the rest of the Maokou Forma-
tion shows uniform δ13Ccarb values at 3.5 ‰, which
is close to the average δ13Ccarb value of the Capit-
anian deposits (c. 4‰, Buggisch et al. 2015) and
records a primary signal. The diagenetic dolostone
succession in the uppermost Maokou Formation is
just below a regional unconformity at the Maokou–
Kuhfeng formations boundary, which is characterized
by a palaeokarst (Fig. 3e, f), probably also indicating a
diagenetic origin of δ13Ccarb at the top of this formation
(e.g. Joachimski, 1994). However, there is no correla-
tion between δ13Ccarb and δ18Ocarb (R2=0.10, Fig. 8a)
in the Wuchiaping Formation, suggesting a primary
signal. The primary δ13Ccarb values in the Wuchiap-
ing Formation are much lower than in the Maokou
Formation.
5.b.2. δ13Corg changes
Organic carbon was drawn from the same dissolved
inorganic carbon (DIC) of carbonate during photo-
synthesis. However, the negative correlation between
δ13Corg and δ13 Ccarb (R2=0.62, Fig. 8b) suggests that
other factors such as organic sources can also con-
trol the δ13Corg changes. The heavy δ13Corg values in
the dolostone succession in the uppermost Maokou
Formation, coincident with the occurrence of bitu-
men (Fig. 7) and the similar values of δ13Corg between
the dolostone succession and the overlying Kuhfeng
chert, suggest that the hydrocarbon bitumen was prob-
ably derived from the Kuhfeng Formation, a hydro-
carbon source rock in South China (Liu et al. 2014),
affected the δ13Corg in the dolostone succession. The
δ13Corg values (c. 22 ‰) of sandstones and claystones
in swamp facies in the lower Wangpo Shale Mem-
ber are much heavier than the δ13Corg of the Kuhfeng
chert in the moderate-water-depth shelf facies. This
suggests an input of terrestrial organic matter sources
in the lower Wangpo Shale Member since marine or-
ganic carbon older than Oligocene age is isotopically
lighter than the land-derived carbon (Galimov, 2006).
Furthermore, the δ13 Corg values in the upper Wangpo
Shale Member and the lower Xiayao Limestone Mem-
ber deposited in a restricted environment near coast-
line are heavier than the δ13Corg values in the upper
Xiayao Limestones deposited in open marine. This
also suggests a more terrestrial 13 C-rich organic mat-
ter input in the former. The input of terrestrial or-
ganic matter therefore controls the δ13Corg changes at
Tianfengping.
5.c. Carbon isotopic changes and their implication for
mass extinction
The carbon-isotope correlation between the Tianfeng-
ping section and the Tieqiao section displays: (1) a
similar trend of δ13Corg between these two sections;
(2) a positive peak at the Maokou–Heshan forma-
tion boundary or the Kuhfeng–Wuchiaping formation
boundary; and (3) similar values of δ13Ccarb in the up-
per Maokou Formation and the lower Wuchiapingian
between these two sections (Fig. 9). These similarit-
ies suggest the δ13Ccarb and δ13 Corg changes represent
regional signals at least since these two sections are
1678 H. WEI AND OTHERS
Figure 9. (Colour online) Carbon-isotope correlation between the Tianfengping section in Hubei Province of South China and the
Tieqiao section in Laibin in Guangxi Province of South China. Organic- and inorganic-carbon isotope data at Tieqiao from Yan, Zhang
& Qiu (2013). Sea-level change at Tieqiao is from Qiu et al.(2014) and Haq & Schutter (2008).
c. 730 km apart from each other. The clastic-origin
tuff/claystones in the lowermost Heshan Formation at
Tieqiao are related to Emeishan volcanism (Zhong,
He & Xu, 2013). The tuff sandstone/claystones in the
Wangpo Shale Member at Tianfengping (Fig. 9) is also
related to the Emeishan volcanism (cf. Isozaki et al.
2008;Heet al. 2010; Deconinck et al. 2014). The
negative shift of δ13Corg values at the G–L boundary
at Tieqiao can be correlated to the same small negat-
ive shift (c. 0.5 ‰in magnitude) of δ13Corg values in
the upper Kuhfeng Formation at Tianfengping. Com-
bined with the conodont zones, the G–L boundary at
Tianfengping is probably in the upper Kuhfeng Form-
ation (Fig. 9). The gradual negative shift of δ13Corg in
the Xiayao Limestone Member in the upper Wuchiap-
ing Formation at Tianfengping can be correlated to
the similar negative shift of δ13Corg in the Clarkina
dukouensis conodont zone at Tieqiao (Fig. 9).
The δ13Ccarb in the lower Wuchiapingian represents
a primary signal and is much lighter (c. 3.0 ‰in mag-
nitude) than that in the lower and middle Maokou
Formation, which also records a primary signal. Previ-
ous studies have suggested that the lower δ13Ccarb val-
ues in the lower Wuchiapingian strata were controlled
by: (1) volcanism and/or thermo-metamorphism meth-
ane (Wignall et al. 2009a;Weiet al. 2012); (2)
low biologic productivity or declined photosynthesis
(Isozaki, Kawahata & Ota, 2007; Yan, Zhang & Qiu,
2013; Nishikane et al. 2014); (3) shallow-marine
anoxia (Saitoh et al. 2013a,2014,2017; Zhang et al.
2015;Weiet al. 2016); and (4) re-oxidation of 12C-
enriched organic material due to eustatic sea-level fall-
ing (Lai et al. 2008). Lithic arenites with abundant
tuff grains in the Wangpo Shale Member in the lower
Wuchiaping Formation at Tianfengping (Fig. 3) sug-
gest a volcanic eruption related to the Emeishan large
igneous province (LIP) (Fig. 5). Generally, large basalt
volcanism releases abundant 12C-enriched CO2into at-
mosphere (Hansen, 2006). However, according to the
model calculation by Berner (2002), the Siberian LIP
can only yield c. 1.0–1.7 ‰magnitude negative excur-
sion in one million years. The much smaller Emeishan
LIP eruption may not yield this large gradual negative
change in δ13Ccarb values in the lower Wuchiapingian
because small volumes of greenhouse gases need to be
released extremely quickly (Jost et al. 2014).
Declined primary productivity decreases the 12C-
enriched organic matter burial, resulting in a neg-
ative excursion of inorganic-carbon isotope (Magar-
itz, 1989; Broecker & Peacock, 1999; Twitchett et al.
2001; Korte & Kozur, 2010). This idea had been ap-
plied to explain the negative excursion of δ13Ccarb
C-isotope and faunal changeover across GLB 1679
across the G–L boundary at Tieqiao by Yan, Zhang
&Qiu(2013). However, the widespread black shale
(Wangpo Shale Member) in the lower Wuchiaping
Formation is enriched in organic matter in the South
China (e.g. at Tianfengping) and North China blocks.
In addition, there were widespread coal beds at the G–
L boundary across the South China Block. The low
value of δ13Ccarb in the lower Wuchiapingian at Tian-
fengping is therefore associated with high organic mat-
ter succession instead of low organic matter, suggest-
ing that low primary productivity may not be the cause
for this negative excursion.
The upwards rising of chemocline permits the return
of isotopically light carbon to shallow-water depths,
resulting in a fall of δ13Ccarb in shallow-water (Küspert,
1982, p. 482; Algeo et al. 2007). The shallow-marine
anoxia had been suggested to be the cause of the neg-
ative excursion of δ13Ccarb across the G–L boundary
Saitoh et al. (2013a). However, the strongest anoxia at
the G–L boundary at Laibin (Wei et al. 2016) is not as-
sociated with the negative excursion of δ13Ccarb (Yan,
Zhang & Qiu, 2013), suggesting that the anoxia may
not have been a major cause for this carbon-isotope
shift.
During global sea-level fall, 12C-enriched organic
matter may have been oxidized at exposed contin-
ental shelves and transported into the ocean (Holser
& Magaritz, 1987,1992; Baud, Magaritz & Holser,
1989). The heavier values of δ13Corg occurring in the
coastal environments instead of the open shallow-
marine environments at Tianfengping indicates that the
input of terrestrial organic matter controls the δ13Corg
changes (e.g. Siegert et al. 2011; Kraus et al. 2013)
because marine organic carbon older than Oligocene
age is isotopically lighter than the land-derived car-
bon (Galimov, 2006) and Permian wood shows heavier
δ13C values than coeval marine-sourced organic mat-
ter (Foster et al. 1997; Krull, 1999; Korte et al. 2001;
Ward et al. 2005; Hermann et al. 2010). The neg-
ative correlation between δ13Ccarb and δ13Corg (R2=
0.62, Fig. 8) therefore suggests that low δ13Ccarb val-
ues correspond to relatively high δ13Corg values dur-
ing/or immediately after regression which brings more
terrestrial organic matter input and results in high
values of δ13Corg at Tianfengping or even in South
China. Large-scale regression during the G–L trans-
ition enhanced the oxidization of exposed organic mat-
ter or soil, resulting in a high input of 12C-rich DIC
via river runoff. In South China, the unconformity at
the G–L boundary is regional (He et al. 2003,2006;
Shen et al. 2007) and represents a regional to global
sea-level fall (Shen et al. 2007; Wignall et al. 2012;
Qiu et al. 2014). This large-scale global regression
(Haq & Schutter, 2008) is associated with the wide-
spread negative excursion of δ13Ccarb in South China
(Wang, Cao & Wang, 2004; Lai et al. 2008; Bond
et al. 2010), Japan (Isozaki, Kawahata & Minoshima,
2007; Isozaki, Kawahata & Ota, 2007) and Iran (Shen
et al. 2013), probably suggesting an impact of regres-
sion on this carbon-isotope change. At Tianfengping,
the Kuhfeng–Wuchiaping formation boundary and the
Maokou–Kuhfeng formation boundary represent re-
gional unconformities and reflect large-scale sea-level
fall. These two pulses of regression may correspond
to the two episodes of Emeishan eruption evidenced
by two basalt successions separated by a siliceous
limestone succession at Xiongjiachang in Guizhou
Province, South China (e.g. Wignall et al. 2009a). Sea-
level fall and re-oxidized organic matter may therefore
be the main controlling factor of this negative shift
of δ13Ccarb in lower Wuchiapingian strata. Qiu et al.
(2013) reported two pulses of carbon isotope excur-
sion during middle Capitanian time and at the G–L
boundary in South China, corresponding to two pulses
of regression. The mechanism of re-oxidized organic
matter can also be used to explain the first pulse of
δ13Ccarb negative excursion during middle Capitanian
time. This sea-level fall coincides with the disappear-
ance of foraminifers and fusulinids at Tianfengping
(Fig. 6), suggesting that sea-level fall may be the cause
of the biotic crisis at the G–L boundary via the loss of
shallow-marine habitat where benthos lived (e.g. Qiu
et al. 2014).
6. Conclusions
The negative shift of δ13Ccarb in the uppermost Maokou
Formation at Tianfengping in South China is of dia-
genetic origin. However, the δ13Ccarb values in the main
part of the Maokou Formation and in the Wuchiap-
ing Formation represent primary signals of coeval sea
water. Bulk organic-carbon isotope changes at Tian-
fengping mainly reflect the shift of organic matter
sources, that is, the increased contribution of terrestrial
organic matter. Sea-level fall led to high terrestrial
organic matter input and resulted in heavy δ13Corg
values. The 3.0 ‰-magnitude lower δ13Ccarb in the
Wuchiaping Formation compared to that in the lower–
middle Maokou Formation at Tianfengping was prob-
ably caused by the re-oxidized 12C-rich organic mat-
ter or soil during sea-level fall. Large-scale global re-
gression resulted in the decrease of δ13Ccarb in lower
Wuchiapingian strata and led to the disappearance of
foraminifers and fusulinids during Capitanian time in
South China.
Acknowledgements. We thank two anonymous reviewers
and the editor Paul Upchurch for their constructive com-
ments and suggestions. This work was financially supported
by the National Natural Science Foundation of China (grant
no. 41302021). We thank Zhijun Zhu’s suggestions for min-
eral identification.
References
Algeo ,T.J.,Ellwood,B.,Nguyen,T.K.T.,Rowe,H.
&May nard, J. B. 2007. The Permian-Triassic bound-
ary at Nhi Tao, Vietnam: Evidence for recurrent in-
flux of sulfidic watermasses to a shallow-marine car-
bonate platform. Palaeogeography, Palaeoclimatology,
Palaeoecology 252, 304–27.
1680 H. WEI AND OTHERS
Ali,J.R.,Thompson, G. M., Zhou,M.F.&Song, X. 2005.
Emeishan large igneous province, SW China. Lithos 79,
475–89.
Baud ,A.,Magaritz,M.&Holser, W. T. 1989. Permian-
Triassic of the Tethys: Carbon isotope studies. Geolo-
gische Rundschau 78, 649–77.
Berner, R. A. 2002. Examination of hypotheses for the
Permo-Triassic boundary extinction by carbon cycle
modeling. Proceedings of the National Academy of Sci-
ences of the United States of America 99, 4172–7.
Bond,D.P.G.,Wignall,P.B.,Joachimski, M. M., Sun,
Y. , S avov ,I.,Grasby,S.E.,Beauchamp,B.&
Blomeier, D. P. G. 2015. An abrupt extinction in the
middle Permian (Capitanian) of the Boreal Realm
(Spitsbergen) and its link to anoxia and acidification.
GSA Bulletin 127, 1411–21.
Bond,D.P.G.,Wignall,P.B.,Wang,W.,Izon,G.,Jiang,
H., Lai,X.,Sun,Y.,Newton,R.J.,Shao,L.,Védrine,
S. & Cope, H. 2010. The mid-Capitanian (Middle Per-
mian) mass extinction and carbon isotope record of
South China. Palaeogeography, Palaeoclimatology, Pa-
laeoecology 292, 282–94.
Broecker,W.S.&Peacock, S. 1999. An ecologic explan-
ation for the Permo-Triassic carbon and sulfur isotope
shifts. Global Biogeochemical Cycles 13, 1167–72.
Buggisch,W.,Krainer,K.,Schaffhauser, M.,
Joachimski,M.&Korte, C. 2015. Late Carbon-
iferous to Late Permian carbon isotope stratigraphy:
A new record from post-Variscan carbonates from
the Southern Alps (Austria and Italy). Pal a e oge o -
graphy, Palaeoclimatology, Palaeoecology 433,
174–90.
Chen,B.,Joachimski,M.M.,Sun,Y.,Shen,S.&Lai,
X. 2011. Carbon and conodont apatite oxygen isotope
records of Guadalupian-Lopingian boundary sections:
Climatic or sea-level signal? Pal aeog eog raph y, Pa l aeo-
climtology, Palaeoecology 311, 145–53.
Chen,L.,Li,Z.,Hua ng,Z.,Zhang,K.&Duan, W. 2000.
Discovery of the clastic rocks in the basal part of
the Permian Gufeng Formation in Huangyan, Jianshi,
Hubei and its significance. Journal of Stratigraphy 24,
207–9 (in Chinese with English abstract).
Chen,W.,Liu,J.,Wang,Z.&Zheng, Q. 2003.
Study on lithofacies palaeogeography during the Per-
mian Emeishan basalt explosion in Guizhou Province.
Journal of Palaeogeography 5, 17–28 (in Chinese with
English abstract).
Chen,Z.Q.,George,A.D.&Yang, W. R. 2009. Effects of
Middle-Late Permian sea-level changes and mass ex-
tinction on the formation of the Tieqiao skeletal mound
in the Laibin area, South China. Australian Journal of
Earth Sciences 56, 745–63.
Clapham,M.E.,Shen,S.&Bottjer, D. J. 2009. The
double mass extinction revisited: reassessing the sever-
ity, selectivity, and causes of the end-Guadalupian biotic
crisis (Late Permian). Paleobiology 35, 32–50.
Courtillot, V. 1999. Evolutionary Catastrophes: The Sci-
ence of Mass Extinction. Cambridge: Cambridge Uni-
versity Press, 237 pp.
Deconinck,J.F.,Craquin,S.,Bruneau,L.,Pellenard,P.,
Baud in,F.&Feng, Q. 2014. Diagenesis of claystone
minerals and K-bentonites in Late Permian/Early Tri-
assic sediments of the Sichuan Basin (Chaotian section,
Central China). Journal of Asian Earth Sciences 81, 28–
37.
Du,X.,Song,X.,Zhang, M., Lu,Y.,Lu,Y.,Chen,P.,Liu,
Z. & Yang, S. 2015. Shale gas potential of the Lower
Permian Gufeng Formation in the western area of the
Lower Yangtze Platform, China. Marine and Petroleum
Geology 67, 526–43.
Feng,Z.,Yang,Y.&Jin, Z. 1997. Lithofacies Paleogeo-
graphy of Permian of South China. Dongying: Petro-
leum University Press, 75 pp. (in Chinese).
Foster,C.B.,Logan,G.A.,Summons,R.E.,Gorter,
J. D. & Edwards, D. S. 1997. Carbon isotopes, kerogen
types and the Permian-Triassic boundary in Australia:
Implications for exploration. Australian Petroleum Pro-
duction and Exploration Association Journal 37, 472–
89.
Galimov, E. M. 2006. Isotope organic geochemistry. Or-
ganic Geochemistry 37, 1200–62.
Hada,S.,Khosithanont,S.,Goto,H.,Fontaine,H.&
Salyapongse, S. 2015. Evolution and extinction of Per-
mian fusulinid fauna in the Khao Tham Yai Limestone
in NE Thailand. Journal of Asian Earth Sciences 104,
175–84.
Hansen, H. J. 2006. Stable isotopes of carbon from basaltic
rocks and their possible relation to atmospheric isotope
excursions. Lithos 92, 105–16.
Haq,B.U.&Schutter, S. R. 2008. A chronology of Paleo-
zoic sea-level changes. Science 322, 64–8.
He,B.,Xu,Y.,Chung,S.,Xiao,L.&Wang, Y. 2003. Sed-
imentary evidence for a rapid, kilometer-scale crustal
doming prior to the eruption of the Emeishan flood
basalts. Earth and Planetary Science Letters 213, 391–
405.
He,B.,Xu,Y.,Wang,Y.&Luo, Z. 2006. Sedimentary and
lithofacies paleogeography in Southwestern China be-
fore and after the Emeishan flood volcanism: New in-
sights into surface response to mantle plume activity.
Journal of Geology 114, 117–32.
He,B.,Xu,Y.G.,Zhong,Y.T.&Guan, J. P. 2010. The
Guadalupian-Lopingian boundary mudstones at Chao-
tian (SW China) are clastic rocks rather than acidic
tuffs: implication for a temporal coincidence between
the end-Guadalupian mass extinction and the Emeishan
volcanism. Lithos 119, 10–19.
Hermann,E.,Hochuli,P.A.,Bucher,H.,Vigran,J.O.,
Weissert,H.&Bernasconi, S. M. 2010. A close-up
view of the Permian-Triassic boundary based on ex-
panded organic carbon isotope records from Norway
(Trøndelag and Finnmark Platform). Global and Plan-
etary Change 74, 156–67.
Holser,W.T.&Magaritz, M. 1987. Events near the
Permian-Triassic boundary. Modern Geology 11, 155–
80.
Holser,W.T.&Magaritz, M. 1992. Cretaceous/Tertiary
and Permian/Triassic boundary events compared.
Geochimica et Cosmochimica Acta 56, 3297–309.
Isozaki, Y. 1997. Permo-Triassic boundary superanoxia and
stratified superocean: Record from lost deep sea. Sci-
ence 276, 235–8.
Isozaki,Y.&Aljinovi´
c, D. 2009. End-Guadalupian extinc-
tion of the Permian gigantic bivalve Alatoconchidae:
end of gigantism in tropical seas by cooling. Pal a eog e o-
graphy, Palaeoclimatology, Palaeoecology 284, 11–21.
Isozaki,Y.,Aljinovic,D.&Kawah ata, H. 2011. The
Guadalupian (Permian) Kamura event in European
Tethys. Palaeogeography, Palaeoclimatology, Palaeoe-
cology 308, 12–21.
Isozaki,Y.,Kawah ata,H.&Minoshima, K. 2007. The
Capitanian (Permian) Kamura cooling event: the begin-
ning of the Paleozoic-Mesozoic transition. Pal a e owo rld
16, 16–30.
Isozaki,Y.,Kawah ata,H.&Ota , A. 2007. A unique car-
bon isotope record across the Guadalupian-Lopingian
C-isotope and faunal changeover across GLB 1681
(Middle-Upper Permian) boundary in mid-oceanic
paleo-atoll carbonates: the high-productivity “Kamura
event” and its collapse in Panthalassa. Global and Plan-
etary Science Change 55, 21–38.
Isozaki,Y.,Yao,J.,Ji,Z.,Saitoh, M., Kobayashi,N.&
Sakai, H. 2008. Rapid sea-level change in the Late
Guadalupian (Permian) on the Tethyan side of Sout
China: litho- and biostratigraphy of the Chaotian sec-
tion in Sichuan. Proceedings of the Japan Academy
Series B 84, 344–53.
Jin,Y.,Shen,S.,Henderson,C.M.,Wang,X.,Wang,
W. , W ang,Y.,Cao,C.&Shang, Q. 2006. The Global
Stratotype Section and Point (GSSP) for the boundary
between the Capitanian and Wuchiapingian Stage (Per-
mian). Episodes 29, 253–62.
Jin,Y.,Zhang,J.&Shang, Q. 1994. Two phases of the
end-Permian mass extinction. Canadian Society of Pet-
roleum Geologists 17, 813–22.
Joachimski, M. M. 1994. Subaerial exposure and deposition
of shallowing upward sequences: evidence from stable
isotopes of Purbeckian pertidal carbonates (basal Creta-
ceous), Swiss and French Jura Mountains. Sedimento-
logy 41, 805–24.
Jost,A.B.,Mundil,R.,He,B.,Brown,S.T.,Altiner,D.,
Sun,Y.,DePaolo,D.J.&Payne, J. L. 2014. Constrain-
ing the cause of the end-Guadalupian extinction with
coupled records of carbon and calcium isotopes. Earth
and Planetary Science Letters 396, 201–12.
Kametaka, M., Takebe, M., Nagai,H.,Zhu,S.&
Takayanagi, Y. 2005. Sedimentary environments of
the Middle Permian phosphorite-chert complex from
the northeastern Yangtze platform, China; the Gufeng
Formation: a continental shelf radiolarian chert. Sedi-
mentary Geology 174, 197–222.
Knauth,L.P.&Kennedy, M. J. 2009. The late Precambrian
greening of the Earth. Nature 460, 728–32.
Kobayashi, F. 2012. Middle and Late Permian foraminifers
from the Chichibu belt, Takachiho area, Kyushu, Japan:
Implications for faunal events. Journal of Paleontology
86, 669–87.
Kofukuda,D.,Isozaki,Y.&Igo, H. 2014. A remark-
able sea-level drop and relevant biotic responses across
the Guadalupian-Lopingian (Permian) boundary in low-
latitude mid-Panthalassa: Irreversible changes recor-
ded in accreted paleo-atoll limestones in Akasaka and
Ishiyama, Japan. Journal of Asian Earth Sciences 82,
47–65.
Korte,C.&Kozur, H. 2010. Carbon-isotope stratigraphy
across the Permian-Triassic boundary: a review. Journal
of Asian Earth Sciences 39, 215–35.
Korte,C.,Veizer,J.,Leythaeuser,D.,Below,R.&
Schwark, L. 2001. Evolution of Permian and Lower
Triassic δ13C in marine and terrigenous organic mater-
ial. Terra N o s t ra 4, 30–4.
Kozur, H. 1993. Upper Permian radiolaria from the
Sosio Valley area, western Sicily (Italy) and from
the uppermost Lamar Limestone of West Texas. Jahr-
buch der Geologischen Bundesanstalt Wien 136(1),
99–123.
Kraus,S.H.,Brandner,R.,Heubeck,C.,Kozur,H.W.,
Struck,U.&Korte, C. 2013. Carbon isotope signa-
tures of latest Permian marine successions of the South-
ern Alps suggest a continental runoff pulse enriched in
land plant material. Fossil Record 16(1), 97–109.
Krull , E. S. 1999. Permian palsa mires as palaeoenviron-
mental proxies. Palaios 14, 530–44.
Küspert, W. 1982. Environmental changes during oil shale
deposition as deduced from stable isotope ratios. In
Cyclic and Event Stratification (eds G. Einsele &
A. Seilacher), pp. 482–501. Berlin: Springer.
Lai,X.,Wang,W.,Wignall,P.B.,Bond,D.P.G.,Jiang,
H., Ali,J.R.,John,E.H.&Sun, Y. 2008. Palaeoen-
vironmental change during the end-Guadalupian (Per-
mian) mass extinction in Sichuan, China. Pal a e oge o -
graphy, Palaeoclimatology, Palaeoecology 269, 78–93.
Liu,G.,Jin,Z.,Luo,K.&Peng, J. 2014. Oil and source cor-
relation in Huangqiao and Jurong areas, Lower Yangtze
region. Petroleum Geology & Experiment 36(3), 359–
66 (in Chinese with English abstract).
Magaritz, M. 1989. 13C minima follow extinction events: a
clued to faunal radiation. Geology 17, 337–40.
Marshall, J. D. 1992. Climatic and oceanographic isotopic
signals from the carbonate rock record and their preser-
vation. Geological Magazine 129, 143–60.
Melim,L.A.,Swart,P.K.&Eberli, G. P. 2004. Mixing-
zone diagenesis in the subsurface of Florida and
the Bahamas. Journal of Sedimentary Research 74,
904–13.
Muttoni,G.,Gaetani, M., Kent,D.,Sciunnach,D.,
Angiolini,L.,Berra,F.,Garzanti,E.,Matte i ,M.
&Zanchi, A. 2009. Opening of the NeoTethys Ocean
and the Pangea B to Pangea A transformation during
the Permian. GeoArabia 14(4), 17–48.
Nishikane,Y.,Kaiho,K.,Henderson, C. M., Takahashi,
S. & Suzuki, N. 2014. Guadalupian-Lopingian con-
odont and carbon isotope stratigraphies of a deep chert
sequence in Japan. Palaeogeography, Palaeoclimato-
logy, Palaeoecology 403, 16–29.
Niu,Z.,Duan,Q.,Fu,T.,Xu,A.,Zeng,B.&Zhu, Y. 2000.
Paleokarst unconformity on top of the Maokou Form-
ation in the Jianshi-Badong area, Hubei: its discovery
and significance. Regional Geology of China 19, 276–
81 (in Chinese with English abstract).
Preto,N.,Spotl,C.&Guaiumi, C. 2009. Evaluation of
bulk carbonate delta C-13 data from Triassic hemipela-
gites and the initial composition of carbonate mud. Sed-
imentology 56, 1329–45.
Qiu,Z.,Wang,Q.,Zou,C.,Yan,D.&Hou, L. 2013.
Carbon isotope negative excursion and its signific-
ance during the Middle–Late Permian transition in the
Laibin area. Geological Review 59(supp.), 1228–31 (in
Chinese).
Qiu,Z.,Wang,Q.,Zou,C.,Yan,D.&Wei, H. 2014.
Transgressive-regressive sequences on the slope of
an isolated carbonate platform (Middle-Late Permian,
Laibin, South China). Facies 60, 327–45.
Rosales,I.,Quesada,S.&Robles, S. 2001. Primary and
diagenetic isotopic signals in fossils and hemipelagic
carbonates: the lower Jurassic of northern Spain. Sedi-
mentology 48, 1149–69.
Saitoh, M., Isozaki,Y.,Ueno,Y.,Yoshida,N.,Yao,J.
&Ji, Z. 2013a. Middle-Upper Permian carbon isotope
stratigraphy at Chaotian, South China: Pre-extinction
multiple upwelling of oxygen-depleted water onto con-
tinental shelf. Journal of Asian Earth Sciences 67–8,
51–62.
Saitoh, M., Isozaki,Y.,Yao,J.,Ji,Z.,Ueno,Y.
&Yoshida, N. 2013b. The appearance of an
oxygen-deplete condition on the Capitanian disphotic
slope/basin in South China: Middle-Upper Permian
stratigraphy at Chaotian. Global and Planetary Change
105, 180–92.
Saitoh, M., Ueno,Y.,Isozaki,Y.,Nishizawa,M.,
Shozugawa,K.,Kawamura,T.,Yao,J.,J
i,Z.,Takai,
K., Yoshida,N.&Mats u o , M. 2014. Isotopic evidence
for water-column denitrification and sulfate reduction
1682 H. WEI AND OTHERS
at the end-Guadalupian (Middle Permian). Global and
Planetary Change 123, 110–20.
Saitoh, M., Ueno,Y.,Matsu ’ura,F.,Kawamura,T.,
Isozaki,Y.,Yao,J.,Ji,Z.&Yoshida, N. 2017. Mul-
tiple sulfur isotope records at the end-Guadalupian (Per-
mian) at Chaotian, China: implications for a role of bi-
oturbation in the Phanerozoic sulfur cycle. Journal of
Asian Earth Sciences 135, 70–9.
Scotese,C.R.&Langford, R. P. 1995. Pangea and paleo-
geography of the Permian. In The Permian of Northern
Pangea (eds P. A. Scholle, T. M. Peryt & D. S. Ulmer-
Scholle), pp. 3–19. Berlin: Springer, vol. 1.
Shen,S.Z.,Cao,C.,Zhang,H.,Bowring,S.A.,
Henderson, C. M., Pay n e ,J.L.,Dav ydo v,V.I.,Chen,
B., Yuan,D.,Zhang,Y.,Wang,W.&Zheng, Q. 2013.
High-resolution δ;13Ccarb chemostratigraphy from latest
Guadalupian through earliest Triassic in South China
and Iran. Earth and Planetary Science Letters 375,
156–65.
Shen,S.Z.&Shi, G. R. 1996. Diversity and extinction pat-
terns of Permian brachiopoda of South China. Histor-
ical Biology 12, 93–110.
Shen,S.Z.&Shi, G. R. 2002. Paleobiogeographical ex-
tinction patterns of Permian brachiopods in the Asian-
western Pacific Region. Paleobiology 28, 449–63.
Shen,S.Z.,Wang,Y.,Henderson,C.M.,Cao,C.&
Wang, W. 2007. Biostratigraphy and lithofacies of the
Permian System in the Laibin-Heshan area of Guangxi,
South China. Pala e owo rld 16, 120–39.
Shi,L.,Feng,Q.,Shen,J.,Ito,T.&Chen, Z. Q. 2016. Pro-
liferation of shallow-water radiolarian coinciding with
enhanced oceanic productivity in reducing conditions
during the Middle Permian, South China: evidence from
the Gufeng Formation of western Hubei Province. Pa-
laeogeography, Palaeoclimatology, Palaeoecology 444,
1–14.
Siegert,S.,Kraus, S.H., Mette,W.,Struck,U.&Korte,
C. 2011. Organic carbon isotope values from the Late
Permian Seis/Siusi succession (Dolomites, Italy): Im-
plications for palaeoenvironmental changes. Fossil Re-
cord 14, 207–17.
Stanley,S.M.&Yang, X. 1994. A double mass ex-
tinction at the end of the Paleozoic Era. Science 266,
1340–4.
Sun,Y.,Lai,X.,Wignall,P.B.,Widdowson,M.,Ali,
J., J iang,H.,Wang,W.,Yan,C.,Bond,D.P.G.&
Védrine, S. 2010. Dating the onset and nature of the
Middle Permian Emeishan large igneous province erup-
tions in SW China using conodont biostratigraphy and
its bearing on mantle plume uplift models. Lithos 119,
20–33.
Twitchett,R.J.,Looy,C.V.,Morante,R.,Visscher,H.
&Wignall, P. B. 2001. Rapid and synchronous col-
lapse of marine and terrestrial ecosystems during the
end-Permian biotic crisis. Geology 29, 351–4.
Wang,W.,Cao,C.&Wang, Y. 2004. The carbon isotope
excursion on GSSP candidate section of Lopingian-
Guadalupian boundary. Earth and Planetary Science
Letters 220, 57–67.
Wang,X.&Sugiyama, T. 2000. Diversity and extinction
patterns of the Permian corals in China. Lethaia 33,
285–94.
Wang,Y.&Jin, Y. 2000. Permian palaeogeographic evol-
ution of the Jiangnan Basin, South China. Pal a e oge o -
graphy, Palaeoclimatology, Palaeoecology 160, 35–44.
Ward,P.D.,Botha,J.,Buick,R.,DeKock,M.O.,Erw i n,
D. H. , Garison,G.H.,Kirschvink,J.L.&Smith,R.
2005. Abrupt and gradual extinction among Late Per-
mian land vertebrates in the Karoo Basin, South Africa.
Science 307, 709–14.
Wei,H.,Chen,D.,Yu,H.&Wang, J. 2012. End-
Guadalupian mass extinction and negative carbon iso-
tope excursion at Xiaojiaba, Guangyuan, Sichuan. Sci-
ence China Earth Science 55, 1480–8.
Wei,H.,Wei,X.,Qiu,Z.,Song,H.&Shi, G. 2016. Redox
conditions across the G-L boundary in South China:
evidence from pyrite morphology and sulfur isotopic
compositions. Chemical Geology 440, 1–14.
Weidlich, O. 2002. Permian reefs re-examined: extrinsic
control mechanisms of gradual and abrupt changes dur-
ing 40 my of reef evolution. Geobios Mémoire Spécial
24, 287–94.
Weissert,H.,Joachimski,M.&Sarnthein, M. 2008. Che-
mostratigraphy. Newsletters on Stratigraphy 42, 145–
79.
Wignall,P.B.,Bond,D.P.G.,Haas,J.,Wang,W.,Jiang,
H., Lai,X.,Altiner,D.,Védrine,S.,Hips,K.,
Zajzon,N.,Sun,Y.&Newton, R. J. 2012. Capitanian
(middle Permian) mass extinction and recovery in west-
ern Tethys: a fossil, facies, and δ13C study from Hun-
gary and Hydra Island (Greece). Palaios 27, 78–89.
Wignall,P.B.,Sun,Y.,Bond,D.P.G.,Izon,G.,Newton,
R. J., Védrine,S.,Widdowson, M., Ali,J.R.,Lai,
X., Jiang,H.,Cope,H.&Bottrell, S. H. 2009a.
Volcanism, mass extinction, and carbon isotope fluc-
tuations in the Middle Permian of China. Science 324,
1179–82.
Wignall,P.B.,Védrine,S.,Bond,D.P.G.,Wang,W.,
Lai,X.,Ali,J.R.&Jiang, H. 2009b. Facies analysis
and sea-level change at the Guadalupian-Lopingian
Global Stratotype (Laibin, South China), and its bear-
ing on the end-Guadalupian mass extinction. Journal of
the Geological Society, London 166, 655–66.
Xia,W.,Zhang,N.,Kakuwa,Y.&Zhang,L.
2006. Radiolarian and conodont biozonation in the
pelagic Guadalupian-Lopingian boundary interval at
Dachongling,Guangxi, South China, and mid-upper
Permian global correlation. Stratigraphy 2, 217–38.
Xu,Y.,Chung,S.,Jahn,B.M.&Wu, G. 2001. Petrolo-
gic and geochemical constraints on the petrogenesis of
Permian-Triassic Emeishan flood basalts in southwest-
ern China. Lithos 58, 145–68.
Yan,D.,Zhang,L.&Qiu, Z. 2013. Carbon and sulfur iso-
topic fluctuations associated with the end-Guadalupian
mass extinction in South China. Gondwana Research
24, 1276–82.
Yao,L.,Gao,Z.,Yang,Z.&Long, H. 2002. Origin of sel-
eniferous cherts in Yutangba Se deposit, southwest En-
shi, Hubei Province. Science China Earth Science 45,
741–54.
Yin,H.F.,Huang ,S.,Zhang,K.,Hansen,H.,Yang,F.,
Ding,M.&Bie, X. 1992. The effects of volcanism
on the Permo-Triassic mass extinction in South China.
In Permo-Triassic Events in the Eastern Tethys (ed. W.
C. Sweet), pp. 146–57. Cambridge: Cambridge Univer-
sity Press.
Yin,H.F.,Jiang,H.,Xia,W.,Feng,Q.,Zhang,N.&Shen,
J. 2014. The end-Permian regression in South China and
its implication on mass extinction. Earth-Science Re-
views 137, 19–33.
Zhang,G.,Zhang,X.,Li,D.,Farquhar,J.,Shen,S.,
Chen,X.&Shen, Y. 2015. Widespread shoaling of
sulfidic waters linked to the end-Guadalupian (Permian)
mass extinction. Geology 43, 1091–4.
Zhang,Z.,Wang,Y.&Zheng, Q. 2015. Middle Permian
smaller foraminifers from the Maokou Formation at the
C-isotope and faunal changeover across GLB 1683
Tieqiao section, Guangxi, South China. Pal aeow o rl d
24, 263–76.
Zhong,Y.T.,He,B.,Mundil,R.&Xu, Y. G. 2014. CA-
TIMS zircon U-Pb dating of felsic ignimbrite from the
Binchuan section: Implications for the termination age
of Emeishan large igneous province. Lithos 204, 14–9.
Zhong,Y.T.,He,B.&Xu, Y. G. 2013. Mineralogy and
geochemistry of claystones from the Guadalupian-
Lopingian boundary at Penglaitan, South China: In-
sights into the pre-Lopingian geological events. Journal
of Asian Earth Sciences 62, 438–62.
Zhou,M.F.,Malpas,J.,Song,X.Y.,Robinson,P.T.,Sun,
M., Kennedy,A.K.,Lesher,C.M.&Keays,R.R.
2002. A temporal link between the Emeishan large ig-
neous province (SW China) and the end-Guadalupian
mass extinction. Earth and Planetary Science Letters
196, 113–22.
Zhu, T. 1989. Sedimentological features and the gen-
esis of lower Permian nodular and thin-bedded sili-
ceous rocks in southern Anhui. Journal of Pa-
laeogeography 43(5), 1–8 (in Chinese with English
abstract).
Content uploaded by Hengye Wei
Author content
All content in this area was uploaded by Hengye Wei on Mar 25, 2019
Content may be subject to copyright.
