www.elsevier.com/locate/palaeo

Palaeogeography, Palaeoclimatology, P

A negative carbon isotope anomaly associated with the earliest

Lopingian (Late Permian) mass extinction

Kunio Kaihoa,*, Zhong-Qiang Chenb, Tomoyuki Ohashia, Tetsuya Arinobuc,

Ken Sawadad, Benjamin S. Cramera

aInstitute of Geology and Paleontology, Tohoku University, Sendai 980-8578, JapanbSchool of Earth and Geographical Sciences, The University of Western Australia, Crawley, WA 3009, Australia

cDepartment of Legal Medicine, Aichi Medical University School of Medicine, Aichi 480-1195, JapandDivision of Earth and Planetary Sciences, Hokkaido University, Sapporo 060-0810, Japan

Received 15 September 2004; received in revised form 24 March 2005; accepted 31 March 2005

Abstract

Palaeozoic life was severely affected in a pair of mass extinction events at the beginning (earliest Lopingian, 260

million years ago) and the end (251 million years ago) of the Late Permian. However, the biological and geochemical

record of the earliest Lopingian crisis remains poorly constrained. We present analyses of strata in South China showing

that a negative anomaly of stable carbon isotope ratios accompanied an abrupt extinction of marine faunas at the beginning

of the Late Permian, similar to patterns observed at the end-Permian and the end-Cretaceous (65 million years ago)

extinction events.

D 2005 Published by Elsevier B.V.

Keywords: Mass extinctions; Permian; C-13/C-12; China

1. Introduction

As the most profound crisis in the history of earth’s

biota, the end-Permian mass extinction triggered the

most widespread reorganization of ecosystems and

animal diversity in the past 500 million years and is

largely responsible for much of the structure of marine

0031-0182/$ - see front matter D 2005 Published by Elsevier B.V.

doi:10.1016/j.palaeo.2005.04.013

* Corresponding author.

E-mail address: kaiho@dges.tohoku.ac.jp (K. Kaiho).

ecosystems today (Bowring et al., 1999). However,

such severe losses among Palaeozoic biota may in fact

have resulted from two extinctions during the Late

Permian (Stanley and Yang, 1994; Jin et al., 1994).

While extensive research has resulted in multidiscipli-

narily evidence documenting the end-Permian crisis

(e.g., Erwin, 1993; Bowring et al., 1998; Jin et al.,

2000; Becker et al., 2001; Kaiho et al., 2001), another

mass mortality taking place in the earliest Lopingian,

~9 million years earlier than the end-Permian event,

has received far less attention. The earliest Lopingian

alaeoecology 223 (2005) 172–180

Fig. 1. Location of the Penglaitan section studied in southern China

K. Kaiho et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 223 (2005) 172–180 173

crisis (50% extinction in marine genera) appears even

stronger in magnitude than the end-Frasnian (41%),

end-Triassic (46%), and end-Cretaceous (45%) mass

extinctions, although it is weaker than the end-Permian

mass extinction (67%) (Sepkoski, 1994). To date, the

earliest Lopingian extinction has received attention

only in South China (Hallam and Wignall, 1997),

where the major benthos (including sponges, corals,

bryozoans, echinoderms, foraminifers, brachiopods,

gastropods, bivalves, and ammonoids) severely suf-

fered from this catastrophe. Although the unique na-

ture of the South China record is likely not typical of

global records (Erwin, 1993), this dramatic biotic loss

has also been reinforced by the latest fossil data of

brachiopods, fusulinids, and corals from South China

(Yang et al., 2000; Wang and Sugiyama, 2000; Shen

and Shi, 2002). However, a stratal record of a discrete

extinction event (as opposed to the stage-level compi-

lation of Sepkoski, 1994) documenting biological and

geochemical perturbations associated with the Guada-

lupian–Lopingian (G/L) crisis has not previously been

presented, although some global (or regional) palaeon-

tologic compilations have documented the biotic loss

at the end of the Guadalupian (Stanley and Yang,

1994; Jin et al., 1994; Yang et al., 2000; Wang and

Sugiyama, 2000; Shen and Shi, 2002). This is the case

because stratigraphically continuous sections encom-

passing the G/L extinction event are rare due to global

regression during the Middle–Late Permian transition

(Ross and Ross, 1987; Hallam and Wignall, 1997).

2. Section studied

This study includes y13C, kerogen, magnetic sus-

ceptibility, and palaeontological analyses of well-pre-

served samples collected from the Penglaitan section

exposed near Laibin, South China (Fig. 1). This strati-

graphic section provides the most continuous G/L

boundary beds in the world (Jin et al., 1998). Here,

the biostratigraphic G/L boundary, which is marked

by the first occurrence of the conodont Clarkina

postbitteri, is placed in the uppermost Maokou For-

mation, ~40 cm below the lithologic contact with the

Heshan Formation (Jin et al., 2001) (Fig. 2). We

obtained unweathered samples by taking large blocks

from the 9-m-thick interval spanning the G/L bound-

ary. A negative excursion of carbonate carbon isotope

.

values across the G/L boundary was reported in this

section (Wang et al., 2004). We present a carbon

isotope record based on y13C values of n�C27 alkane

and bulk carbonate that is more detailed in the Heshan

Formation and shows a different pattern of variations

in the Maokou Formation.

3. Methods

3.1. Magnetic susceptibility

Magnetic susceptibility was measured on the out-

crop using a Bartington MS/2 magnetic susceptibility

meter with a MS2F probe capable of obtaining inde-

pendent measurements at 2 cm intervals. Reported

values are not corrected for volume.

3.2. Carbon isotope

Prior to analyses of the carbon isotope ratio of

alkane, the surfaces of rock chip samples were

cleaned three times by ultrasonication using a ben-

zene/methanol solvent mixture (6 :4, v/v). Each sam-

ple (200 g) was crushed and ground into powder and

extracted three times with a benzene/methanol sol-

vent mixture (6 :4, v/v) by ultrasonic agitation. The

-1

-2

-3

-4

5

4

2

1

0

3

Up

pe

r P

erm

ian

Gu

ad

alu

pia

nLo

ping

ian

Mao

kou

Hes

han

Thi

ckne

ss (

m)

Pen

glai

tan

Sec

tion

Ser

ies

For

mat

ion

Crinoid limestone Limestone

Limestone nodule

Marl

Muddy limestone nodule

Benthic fossil fragments (%)

Dec

reas

e in

ben

thic

foss

il fr

agm

ents

Last

occ

urre

nce

ofG

uada

lupi

an ta

xa

-29 -27

13C27 n-alkane (‰)

3 4 513C carbonate

(‰)

-28

Siliceous mudstone

30

Bed

num

ber

B

12

0

29

21

7

14

2019

18

22232425

-1

-2

-4

-5

-7

-6

-8

Tuff

-9-10-11

-13

-14

-16

-19

Sta

ge

Bed

num

ber

A

6

5

4

3c

k

k

a

e

aMid

dle

P

erm

ian

Clay

20 40 60 80 40 80

Marine kerogens (%)

-26 -18-22

13C organic (‰)

Sponge

BrachiopodaCrinoid

Coral

Gastropoda

Foraminifera

Bryozoa

Total of benthic fossils

δ δ

δ

Fig. 2. Variations in lithology, benthic faunal groups based on fragments, marine kerogens, and carbon isotopes recorded in Middle and Upper

Permian beds in Penglaitan section, southern China. Benthic fossil fragments (%) is the % of a thin section’s area occupied by the fossils.

Marine kerogens (%) is percentage of marine kerogens/marine plus terrestrial kerogens. Bed number A is after Henderson et al. (2002). Bed

number B is reported in this paper. Definition and placement of the Guadalupian–Lopingian boundary is after Henderson et al. (2002). PDB is

the Peedee belemnite. Samples from beds 9 and 21 were taken from limestone nodules. The last occurrence of Guadalupian taxa include corals

(Ufimia), brachiopods (Uncisteges), fusulinids (Yabeina, Metadoliolina, Chusenella, Kahlerina, Lantschichites, Lepidolina), and conodonts

(Jinogondolella) (Jin et al., 1998, 2001). The broken lines connecting points from the two data sets of y13C were drawn to suggest that the

combined dataset indicates a change of y13C in the atmosphere–ocean carbon reservoir.

K. Kaiho et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 223 (2005) 172–180174

extract was fractionated into neutral and acidic com-

pounds by liquid–liquid separation. The neutrals

were fractionated into aliphatic hydrocarbons and

other classes of compounds by silica-gel column

chromatography (Silica gel 1% water; column size:

50�5 mm i.d.). A molecular sieve (type 5A) was

used to isolate individual n-alkanes from the aliphat-

ic hydrocarbon fraction. The y13C of n-alkanes was

measured using a Finnigan Mat GC/IRMS (gas chro-

matography/isotope ratio mass spectrometer) system.

The GC oven temperature for n-alkanes was

programmed from 50 to 280 8C at ramp rate of 30

8C/min, and then increased to 310 8C at 5 8C/min.

The overall uncertainties of y13C measurements are

within 0.4x. The carbon isotope ratio of bulk car-

bonate sediments were analyzed using a Finnigan

Mat Delta-S mass spectrometer. The overall uncer-

tainties of y13C are within 0.2x. The carbon isotope

ratio of bulk organic carbon were analyzed using a

Finnigan Mat Delta Plus Advantage mass spectrom-

eter (Continuous Flow Isotope Ratio MS). The over-

all uncertainties of y13C are within 0.3x. All

K. Kaiho et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 223 (2005) 172–180 175

isotopic values are reported relative to the Peedee

belemnite (PDB) standard.

3.3. Fossil

The quantitative point count data (Fig. 2) are

based on detailed sample sections focusing on the

microfacies and palaeontology of the G–L boundary

beds in Laibin. The point-counting data from each

thin-section are based on observation of 20 to 50

views in one sample. Percentages of various skeletal

components, micrite, silica and cavities were com-

bined to yield the mean abundance of each compo-

nent for the sample.

3.4. Kerogen

Whole rock samples were crushed to a fine powder

in an agate mortar. Powdered rock samples (5 to 10 g)

were treated sequentially within water bath shaker as

follows: HCl 6 M (100 ml, 60 8C, 12 h), HCl 12 M/

HF 46% (1 /1 v/v) (100 ml, 60 8C, and 24 h), HCl 6 M(100 ml, 60 8C, 4 h). After each treatment, the super-

natant was removed after centrifugation (3000 rpm, 10

min). The residue, kerogen, was sequentially washed

with HCl 6 M (�2) and distilled water (�5), and was

recovered and freeze-dried under vacuum.

Wet residual particles obtained after HCl/HF

treatment were siphoned and dropped to slides,

and were mounted in permanent slides for micro-

scopic examination. Three hundred particles of ker-

ogen macerals were point counted per slide under

transmitted light microscope. Fluorescence light mi-

croscopic observation was used to differentiate be-

tween marine-derived and terrestrially derived

amorphous kerogens since the former tends to fluo-

resce significantly (Sawada and Akiyama, 1994).

However, this method should be used with caution,

because the terrigenous amorphous kerogen formed

by the advanced degradation of herbaceous macerals,

such as sporinite and cutinite, might fluoresce (Smith,

1984; Oboh-Ikuenobe et al., 1997) and mature amor-

phous kerogen does not fluoresce even when derived

from marine organisms. Therefore, we examined the

identification of amorphous kerogens on the basis of

both the fluorescence and the morphology under

transmitted light microscope as reported by Boussafir

et al. (1995).

4. Results and discussion

4.1. Description of section

In Penglaitan, biostratigraphic and sedimentologic

studies (Jin et al., 1998; Mei et al., 1994, 1998) reveal

that the top 9 m thick limestone of the Maokou

Formation is a shallowing-up succession reflecting

the global fall in sea-level at the end of the Middle

Permian (Ross and Ross, 1987; Chen et al., 1998).

The basal Heshan Formation is an upward-deepening

succession reflecting a regional transgression in the

Tethyan region (Chen et al., 1998). Field observations

of probable Milankovitch-driven lithologic variations

[see Appendix] suggest that sedimentation rates may

have been 11.0–13.5 cm/k.y. in the uppermost part of

the Maokou Formation. Lithologic cyclicity in the

lowermost Heshan Formation can plausibly be asso-

ciated with precessional forcing, which would imply

an abrupt decrease in sedimentation rates to 3.3–4.1

cm/k.y. immediately above the boundary with a sub-

sequent increase (to 4.8–5.9 cm/k.y. by bed 18). Al-

though future temporal constraints may lead to a

revision of this orbital stratigraphy, we use it here as

the best available estimate of sedimentation rates in

this section and as the basis for estimates of rates and

durations given below.

4.2. Abrupt mass extinctions

The continuous succession of the Penglaitan sec-

tion provides a unique opportunity to investigate the

stratal record of the G/L mass extinction, which has

previously only been inferred from statistical data.

Although few macrofossils have been described

from this section, we found the Laibin Limestone

(uppermost Maokou Formation) to be exceptionally

fossiliferous with common to abundant brachiopods,

corals, ammonoids, and crinoids visible on the expo-

sures. In contrast, the overlying Heshan Formation is

poorly fossiliferous except for a few ammonoids and

trace fossils. Of particular significance is the last

appearance of numerous index fossils close to the

G/L boundary (Fig. 2; Jin et al., 1998). This pattern

suggests an abrupt biotic loss slightly after the begin-

ning of the Lopingian (Fig. 2). The pattern of fossil

extinctions is strengthened by the stratigraphic distri-

bution of benthic fossil fragments counted in thin

K. Kaiho et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 223 (2005) 172–180176

sections (Fig. 2). We found that the abundance of

fossil fragments of sponges, bryozoans, brachiopods,

crinoids, foraminifers, corals, and gastropods de-

creased from 86% to 0% in the 18 cm interval across

the Maokou/Heshan Formation (Fig. 2). These data

suggest that a mass extinction occurred within a short

temporal duration corresponding to the uppermost 18

cm of the Maokou Formation (~1.5 k.y.; our estimate

of such a short duration is obviously subject to large

error but is probably correct within a factor of 10), in

the lowermost Lopingian.

Because the rapid decrease in fossil diversity cor-

responds to a lithologic transition from limestone

(Maokou Formation) to siliceous mudstone (Heshan

Formation), it is appropriate to consider the possibility

that the apparent extinction may simply reflect a local

facies shift. Our analyses of kerogens indicate that the

percentage of marine kerogens/marine plus terrestrial

kerogens does not change significantly across the

extinction boundary (Fig. 2; 62 to 78% in beds 0

and 1 and 50 to 98% between beds �13 and 9).

This implies that the rapid decrease in fossils occurred

without any significant change in the proportion of

marine vs terrestrial kerogens. It is therefore unlikely

that the disappearance of calcareous fossils at the top

of the Maokou Formation was due to an increase in

the input of terrestrial material associated with the

facies shift. Instead, the extinction of calcareous

biota and the associated environmental perturbation

most likely caused the lithologic change from lime-

stone to mudstone, similar to facies changes often

observed in records of the end Permian and K/T

boundary events. We conclude that the sharp decline

in biotic abundance and diversity in the Penglaitan

section most likely signals the G/L boundary mass

extinction previously inferred from statistic palaeon-

tologic data.

4.3. Negative shift of carbon isotope values

Carbon isotope ratios of long-chain n-alkane

(C26–C30), which are derived from terrestrial higher

plant wax, do not vary with maturity (Ishiwatari et

al., 1999). Therefore, y13C variations of long-chain

n-alkane remaining in older sediments represent ma-

terial derived from terrestrial higher plant wax. We

measured the y13C of higher molecular weight

n�C27 alkane, and the range of y13C variation

was approximately �29x to �27x (Fig. 2), similar

to y13C of terrestrial higher plants. The y13C of bulk

organic carbon shows a much larger range between

approximately �27x to �19x (Fig. 2), which we

take to be due to variations in the ratio of marine and

terrestrial organic carbon content superimposed on

variations in the y13C of atmospheric CO2. The y13Cvalues of n�C27 alkane are lower than those of bulk

organic carbon, which supports the interpretation of

y13C values of n�C27 alkane as representing mate-

rial derived from terrestrial higher plants. For our

interpretation, we have assumed that the y13C shift

for the n�C27 alkane is a direct reflection of a shift

in y13C values of atmospheric CO2 (Arinobu et al.,

1999).

A 1.1x negative shift in y13C values of n�C27

alkane occurred between beds 1 and 3 (bed number in

the Penglaitan Section) (Fig. 2). Values of y13C grad-

ually decrease by 1.1x from �26.9x to �28.0x in

the top 4 m of the Maokou Formation; followed by an

abrupt 1.1x decrease to �29.1x in the basal part (50

cm) of the Heshan Formation. Values then gradually

return by 1.5x to �27.7x at 2 m above the mini-

mum value. Although carbon isotope ratios of marine

bulk carbonate are sporadic in the basal 2 m of the

Heshan Formation, the y13C of marine bulk carbonate

shows mostly synchronized changes with that of

n�C27 alkane ranging between 3.2x and 4.8x(Fig. 2).

The y13C values of n�C27 alkane and those of

carbonate represent the y13C of the global surface

reservoir (atmosphere, surface ocean, and biota).

The compilation of both data sets (Fig. 2) indicates

that an abrupt negative excursion of ~2x occurred

between beds 2 and 3, some 46 cm (11–14 k.y.) above

the decrease in benthic fossil fragments. The rapidity

of the 2x decrease in y13C values, which occurred

between two samples, cannot be resolved in our re-

cord and likely occurred in ~1 k.y. or less. The

recovery occurred within 2 precession cycles (34–42

k.y.) following the minimum value (This follows the

description in the paragraphs above, indicating the

return to pre-excursion values 2 m above the mini-

mum value; if recovery is instead taken at the values

~28x near bed 7, then the duration would be a single

precession cycle, or ~17–21 k.y.). The y13C anomaly

is therefore similar in pattern and duration to that at

the end-Cretaceous mass extinction: an abrupt major

Upp

er

Per

mia

n

Gua

dalu

pian

Lopi

ngia

n

Mao

kou

Hes

han

Ser

ies

For

mat

ion

Sta

geM

iddl

e P

erm

ian

Tert

iary

Cre

t.P

erio

d

Thi

ckne

ss (

m)

δ C carbonate vs PDB (‰)13

δ C n-alkane 13

27vs PDB (‰)

δ C n-alkane 13

29vs PDB (‰)

Tim

e (k

yr)

-10

0

10

20

30-29 -28 -27

-29 -28 -27

-4

-3

-2

-1

0

1

2

3

4

3 4 5

Fig. 3. Correlation between variations in y13C across the Guadalu-

pian–Lopingian boundary in the Penglaitan section, southern China

and similar trends recorded across the Cretaceous–Tertiary bound-

ary in Caravaca, Spain (Arinobu et al., 1999). PDB is the Peedee

belemnite. Arrows indicate the mass extinction horizon.

K. Kaiho et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 223 (2005) 172–180 177

decrease in b1 k.y. follows a more gradual increase

during several 10s of k.y. (Fig. 3).

5. Conclusions

The pattern, shown here, that a rapid decrease in

biotic diversity is followed by an abrupt negative

anomaly in carbon isotope values is comparable

with the record of the end-Permian mass extinction

in the Meishan section, South China (Bowring et al.,

1998; Jin et al., 2000; Kaiho et al., 2001). Our results

indicate that the G/L boundary mass extinction took

place in a single abrupt event in the earliest Lopin-

gian, corresponding to bed 1 in the basal Heshan

Formation in Penglaitan, South China.

Several studies (Hallam and Wignall, 1997, 1999;

Shi et al., 1999) have considered that Late Guadalu-

pian global regression was principally accountable for

the G/L mass extinction. Alternatively, Isozaki (1997)

suggested that the onset and climax of global oceanic

stratification across the Permian/Triassic (P/T) bound-

ary correspond with the G/L and end-Permian extinc-

tions, respectively. Courtillot et al. (1999) and Zhou et

al. (2002) proposed that the Emeishan flood basalts

(600 km diameter) in South China were erupted at the

G/L Transition, causing the extinction. However, our

study demonstrates that the dramatic losses of biota

coincide with the extremely negative shift of carbon

isotope curves. The comparable patterns of biotic

extinction and negative shift in y13C values suggest

that both the earliest Lopingian and end-Permian mass

extinctions were catastrophic events and may have

resulted from similar causes, although the direct cat-

astrophic causes remain unresolved at present.

Acknowledgments

This work was partly supported by a grant-in-aid

for scientific research from the Japan Society for the

Promotion of Science. This work was conducted as

part of the 21st Center-Of-Excellence program,

dAdvanced Science and Technology Center for the

Dynamic EarthT of Tohoku University. We thank D.

H. Erwin, P. B. Wignall, K.-U. Hinrichs, and J. C.

Ingle for comments.

Appendix A. Orbital cyclicity

In the upper Maokou Formation we observed an

alternation between mainly thick-bedded and mainly

thin-bedded chert/limestone intervals. These packages

are in turn a part of a repetitive alternation between

more easily erodable packages of (alternating thin-and

thick-bedded) chert/limestone and more massive

muddy limestones with mainly thick beds. We believe

that these repetitive lithologic variations reflect pre-

cessional forcing (resulting in alternating thick-and

thin-bedded intervals) modulated by the ~100 k.y.-

period eccentricity component (resulting in alternating

bedded and massive packages). On a larger scale,

there is a cyclic variation in the degree to which the

massive intervals project into the river that is consis-

tent with modulation by the ~400 k.y.-period long

Fig. 4. Two photographs of the lowermost Heshan Formation with the magnetic susceptibility record measured on the outcrop. Arrows indicate

thicker beds (1, 7, and 18 in bed number B scheme), which are associated with low magnetic susceptibility. Note that the vertical scaling of the

outcrop in the image is non-linear, so that the magnetic susceptibility log does not precisely correspond to the image.

K. Kaiho et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 223 (2005) 172–180178

eccentricity cycle. Similar bundling of cycles consis-

tent with orbital forcing has been used to construct

orbital timescales for sections representing a wide

variety of depositional environments throughout the

Mesozoic and Cenozoic (e.g., Herbert and D’Hondt,

1990; Hilgen, 1991; Olsen and Kent, 1996; Rampino

et al., 2000). The lithologic variation is not as easily

discernible in the Heshan Formation, but we did

observe repetitive topographic highs projecting into

the river, with a similar repetition interval to those in

the Maokou Formation, and in some places (including

immediately above the boundary) a cyclic variation in

bed thickness is evident (Fig. 4).

We associate the massive intervals in the Maokou

Formation with eccentricity minima, of which bed 3

(Bed number A scheme; Fig. 2) is the uppermost. This

limits the duration represented by the uppermost ~6 m

of the Maokou Formation to b100 k.y. We further

note, though, that beds 4 and 6 are more thinly bedded

while bed 5 is more thickly bedded, and therefore

suggest that the ~2.3 m interval from the top of bed

3 to the top of bed 5 represents a single precession

cycle. Assuming a duration of 17–21 k.y. for individ-

ual precession cycles at this time (Berger et al., 1992),

this leads to a sedimentation rate estimate for this

interval of 11.0–13.5 cm/k.y. It is also plausible that

the cyclic variation in the lowermost Heshan Forma-

tion from generally thick-bedded to thin-bedded and

back to thickly bedded (also reflected in low-high-low

magnetic susceptibility cycles; Fig. 4) reflects preces-

sional forcing of insolation. If so, then the sedimen-

tation rate in the ~0.7 m interval from the base of bed

1 (Bed number B scheme) to the base of bed 7 is in

the range 3.3–4.1 cm/k.y. while in the ~1.0 m interval

from the base of bed 7 to the base of bed 18 it is in the

range 4.8–5.9 cm/k.y. We emphasize that interpreta-

tion of individual precession cycles in intervals of

climatic and lithologic perturbations should always

be regarded as speculative, since such perturbations

are likely to result in large variations in sedimentation

rates and the climate response to insolation forcing as

well as possibly cyclic climatic variations unrelated to

orbital forcing.

References

Arinobu, T., Ishiwatari, R., Kaiho, K., Lamolda, M.A., 1999. Spike

of polycyclic aromatic hydrocarbons associated with an abrupt

decrease in y13C of a terrestrial biomarker at the Cretaceous–

Tertiary boundary at Caravaca, Spain. Geology 27, 723–726.

Becker, L., Poreda, R.J., Hunt, A.G., Bunch, T.E., Rampino, M.,

2001. Impact event at the Permian–Triassic boundary: evidence

from extraterrestrial noble gases in fullerenes. Science 291,

1530–1533.

Berger, A., Loutre, M.F., Laskar, J., 1992. Stability of the astro-

nomical frequencies over the Earth’s history for paleoclimate

studies. Science 255, 560–566.

K. Kaiho et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 223 (2005) 172–180 179

Boussafir, M., Gelin, F., Lallier-Verges, E., Derenne, S., Bertrand,

P., Largeau, C., 1995. Electron microscopy and pyrolysis of

kerogens from the Kimmeridge Clay Formation, UK: source

organisms, preservation processes, and origin of microcycles.

Geochimica et Cosmochimica Acta 59, 3731–3747.

Bowring, S.A., Erwin, D.H., Jin, Y.G., Martin, M.W., David, E.K.,

Wang, W., 1998. U/Pb zircon geochronology and tempo of the

end-Permian mass extinction. Science 280, 1039–1045.

Bowring, S.A., Erwin, D.H., Isozaki, Y., 1999. The tempo of mass

extinction and recovery: the end-Permian example. Proceedings

of the National Academy of Sciences of the United States of

America 96, 8827–8828.

Chen, Z.Q., Jin, Y.G., Shi, G.R., 1998. Permian transgres-

sion–regression sequences and sea-level changes of

South China. Proceedings of the Royal Society of Victoria

110, 345–367.

Courtillot, V., Jaupart, C., Manighetti, I., Tapponnier, P., Besse,

J., 1999. On causal links between flood basalts and con-

tinental breakup. Earth and Planetary Science Letters 166,

177–195.

Erwin, D.H., 1993. The Great Paleozoic Crisis. Columbia Univer-

sity Press, New York. 327 pp.

Hallam, A., Wignall, P.B., 1997. Mass Extinctions and their After-

math. Oxford University Press, Oxford. 320 pp.

Hallam, A., Wignall, P.B., 1999. Mass extinctions and sea level

changes. Earth-Science Reviews 48, 217–250.

Henderson, C.M., Mei, S.L., Wardlaw, B.R., 2002. New conodont

definitions at the Guadalupian–Lopingian boundary. In: Hills,

L.V., Henderson, C.M., Bamber, E.W. (Eds.), Carboniferous and

Permian of the World, Memoir-Canadian Society of Petrolium

Geologists, vol. 19, pp. 725–735.

Herbert, T.D., D’Hondt, S.L., 1990. Precessional climate cyclicity in

Late Cretaceous–Early Tertiary marine sediments: a high reso-

lution chronometer of Cretaceous–Tertiary boundary events.

Earth and Planetary Science Letters 99, 263–275.

Hilgen, F.J., 1991. Astronomical calibration of Gauss to Matuyama

sapropels in the Mediterranean and implication for the Geomag-

netic Polarity Time Scale. Earth and Planetary Science Letters

104, 226–244.

Ishiwatari, R., Kon, M., Yamada, K., Arinobu, T., Naraoka, H.,

Okui, A., Waseda, A., 1999. Application of gas chromatog-

raphy/isotope ratio mass spectorometry to petroleum geo-

chemistry: carbon isotopic composition of individual long

chain normal alkanes in oils from Northeast Japan. Journal

of the Japanese Association of Petroleum Technologist 64,

551–559.

Isozaki, Y., 1997. Permo-Triassic boundary superanoxia and strat-

ified superocean: records from lost deep sea. Science 276,

235–238.

Jin, Y.G., Zhang, J., Shang, Q.H., 1994. Two phases of the end-

Permian mass extinction. Memoir-Canadian Society of Petro-

leum Geologists 17, 813–822.

Jin, Y.G., Mei, S.L., Wang, W., Wang, X.D., Shen, S.Z., Shang,

Q.H., Chen, Z.Q., 1998. On the Lopingian series of the Permian

system. Palaeoworld 9, 1–18.

Jin, Y.G., Wang, Y., Wang, W., Shang, Q.H., Cao, C.Q., Erwin,

D.H., 2000. Pattern of marine mass extinction near the

Permian–Triassic boundary in south China. Science 289,

432–436.

Jin, Y.G., Henderson, C.M., Wardlaw, B.R., Glenister, B.F., Mei,

S.L., Shen, S.Z., Wang, X.D., 2001. Proposal for the Global

Stratotype Section and Point (GSSP) for the Guadalupian–

Lopingian Boundary. Permophiles 39, 32–42.

Kaiho, K., Kajiwara, Y., Nakano, T., Miura, Y., Kawahata, H.,

Tazaki, K., Ueshima, M., Chen, Z., Shi, G.R., 2001. End-

Permian catastrophe by a bolide impact: evidence of a

gigantic release of sulfur from the mantle. Geology 29,

815–818.

Mei, S.L., Jin, Y.G., Wardlaw, B.R., 1994. Zonation of conodonts

from the Maokouan–Wuchiapingian boundary strata, South

China. Palaeoworld 4, 225–234.

Mei, S.L., Jin, Y.G., Wardlaw, B.R., 1998. Conodont succes-

sion of the Guadalupian–Lopingian boundary strata in Lai-

bin of Guangxi, China and West Texas, USA. Palaeoworld

9, 53–76.

Oboh-Ikuenobe, F.E., Yepes, O., ODP Leg 159 scientific

party, 1997. Palynofacies analysis of sedimentary from

the Cote d’Ivoire-Ghana transform margin: preliminary

correlation with some regional events in the Equatorial

Atlantic. Palaeogeography, Palaeoclimatology, Palaeoecology

29, 291–314.

Olsen, P.E., Kent, D.V., 1996. Milankovitch climate forc-

ing in the tropics of Pangaea during the Late Tri-

assic. Palaeogeography, Palaeoclimatology, Palaeoecology 122,

1–26.

Rampino, M.R., Prokoph, A., Adler, A., 2000. Tempo of the end-

Permian event: high-resolution cyclostratigraphy at the Perm-

ian–Triassic boundary. Geology 28, 643–646.

Ross, C.A., Ross, J.R.P., 1987. Late Paleozoic sea-levels and de-

positional sequence. In: Ross, C.A., Haman, D. (Eds.), Timing

and Depositional History of Eustatic Sequences: Constraints on

Seismic Stratigraphy. Cushman Foundation for Foraminiferal

Research Special Publication, vol. 24, pp. 137–149.

Sawada, K., Akiyama, M., 1994. Carbon isotope composition of

macerals separated from various kerogens by density separation

method. Journal of the Japanese Association of Petroleum Tech-

nologist 59, 244–255 (in Japanese).

Sepkoski Jr., J.J., 1994. Extinction and the fossil record. Geotimes,

15–17 (March).

Shen, S.Z., Shi, G.R., 2002. Paleobiogeographical extinction pat-

terns of Permian brachiopods in the Asian–western Pacific

region. Paleobiology 28, 449–463.

Shi, G.R., Shen, S.Z., Tong, J.N., 1999. Two discrete, possibly

unconnected, Permian marine mass extinctions. In: Yin, H.F.,

Tong, J.N. (Eds.), Proceedings of the International Conference

on Pangea and the Paleozoic–Mesozoic transition. China Uni-

versity of Geosciences Press, Wuhan, pp. 148–151.

Smith, P.M.R., 1984. The use of fluorescence microscopy in the

characterization of amorphous organic matter. Organic Geo-

chemistry 6, 839–845.

Stanley, S.M., Yang, X.N., 1994. A double mass extinction at the

end of the Paleozoic Era. Science 266, 1340–1344.

Wang, X.D., Sugiyama, T., 2000. Diversity and extinction patterns

of Permian coral faunas of China. Lethaia 33, 285–294.

K. Kaiho et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 223 (2005) 172–180180

Wang, W., Cao, C., Wang, Y., 2004. The carbon isotope excursion

on GSSP candidate section of Lopingian–Guadalupian bound-

ary. Earth and Planetary Science Letters 220, 57–67.

Yang, X.N., Shi, G.J., Liu, J.R., Chen, Y.T., Zhou, J., 2000. Inter-

taxa differences in extinction process of Maokouan (Middle

Permian) fusulinaceans. Science in China, Series D: Earth

Sciences 43, 633–637.

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 igneous province (SW China)

and the end-Guadalupian mass extinction. Earth and Planetary

Science Letters 196, 113–122.