Palaeogeography, Palaeoclimatology, Palaeoecology 274 (2009) 32–39
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Palaeogeography, Palaeoclimatology, Palaeoecology
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / p a l a e o
Carbon and sulfur isotopic anomalies across the Ordovician–Silurian boundary on the
Yangtze Platform, South China
Detian Yan a,c, Daizhao Chen b,⁎, Qingchen Wang c, Jianguo Wang b, Zhuozhuo Wang b
a
b
c
Faculty of Earth Resources, China University of Geosciences, Wuhan 430074, China
Key Laboratory of Petroleum Geophysics, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China
State Key Laboratory of Lithosphere Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China
a r t i c l e
i n f o
Article history:
Received 27 February 2008
Received in revised form 22 November 2008
Accepted 23 December 2008
Keywords:
Stable isotopes
Carbon–sulfur isotope
Mass extinctions
Climate change
Anoxia
Ordovician–Silurian boundary
Yangtze Platform
a b s t r a c t
Abundance of organic matter, carbon isotopic compositions of organic matter (δ13Corg) and sulfur isotopic
compositions of iron sulfide (δ34Ssulfide) across the Ordovician–Silurian (O–S) boundary was analysed from
two sections (Wangjiawan, western Hubei; Nanbazi, northern Guizhou) on the Yangtze Platform, South
China. The organic abundance across the O–S boundary at these two sections is generally high, except for the
Hirnantian interval. At Wangjiawan section in western Hubei, the δ13Corg values vary from about −30.8‰
VPDB in the mid-Ashgill to −27.6‰ in the upper Hirnantian interval, which abruptly return to the preHirnantian values. At Nanbazi section in northern Guizhou, δ13Corg values vary from − 30.5‰ to −26.6‰ from
the mid-Ashgill to the upper Hirnantian horizons, which shift negatively to the low spike at −29.2‰ in the
lowermost Rhuddanian, then increase slightly to relatively persistent values around −28‰. Similar variation
patterns of δ34Ssulfide values are unravelled at the two sections, showing a positive excursion from the midAshgill to the upper Hirnantian, then a sharp negative shift in the lowermost Rhuddanian, although with
different background values of δ34Ssulfide from these two sections. These data, together with those from other
areas, demonstrate large climatic fluctuations from warming to cooling, then to warming, oceanic changes
from and anoxic to oxygenated, to anoxic water columns, during the O–S transition. Climatic fluctuations,
together with multiple oceanic anoxia and sea-level fluctuations, were likely responsible for the stepwise
massive demise of the O–S biotic crisis.
© 2009 Elsevier B.V. All rights reserved.
1. Introduction
The Ordovician–Silurian (O–S) transition was a critical interval in
geological history, during which profound biotic, climatic and tectonic
changes occurred (Brenchley et al., 1994, 2003; Kump et al., 1999;
Wang et al., 1997; Finney et al., 1999; Chen et al., 2004); two phases of
extinction events, coincided with the major sea-level changes, were
identified (Brenchley, 1988). Several hypotheses, including bolide
impact (Wilde et al., 1986; Goodfellow et al., 1992), productivity
collapse (Brenchley et al., 1994) and weathering processes (Kump et al.,
1999), have been proposed to be responsible for this particular geologic
event. The ultimate cause, however, is still highly controversial.
Extensive studies revealed a large-magnitude positive carbon
excursion (up to 7‰ PDB locally) in the Hirnantian stage (N.
extraordinarius–N. ojsuensis zone) of the latest Ordovician (e.g.,
Brenchley et al., 1994, 2003; Marshall et al., 1997; Finney et al.,
1999; Kump et al., 1999; Kaljo et al., 2004). An organic-carbon isotopic
excursion of similar to somewhat larger magnitudes has been
⁎ Corresponding author.
E-mail address: [email protected] (D. Chen).
0031-0182/$ – see front matter © 2009 Elsevier B.V. All rights reserved.
doi:10.1016/j.palaeo.2008.12.016
corroborated in the Hirnantian shales as well (e.g., Melchin and
Holmden, 2006). By analogy to carbon cycles, sulfur reservoir sizes
and redox variations within the sedimentary sulfur cycle account for
the observed variations in δ34S values (Burdett et al., 1989; Strauss,
1997). Until now, a few comparable sulfur isotopic studies across
several critical boundaries were conducted (Geldsetzer et al., 1987;
Kajiwara et al., 1994; Joachimski et al., 2001; Newton et al., 2004; Fike,
2007; Gill et al., 2007; Riccardi et al., 2007). In this paper, we present
new carbon and sulfur isotopic chemostratigraphic data across the O–
S boundary, which provide useful clues to understand the tremendous
climatic, oceanic changes and relevant biotic crisis during this critical
transition.
2. Geological setting
South China block appears to have been operated as a distinct plate
during the mid-Paleozoic, but it was still attached to the margin of
Gondwana during the Late Ordovician to Early Silurian (Fig. 1A)
(Metcalfe, 1994). One of the most important features of this tectonic
unit in the Ordovician was that the Yangtze Platform, covered by a
broad epeiric sea, was bordered southeast by the deep Pearl River Sea,
such that it was likely connected more or less to the open ocean
D. Yan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 274 (2009) 32–39
33
Fig. 1. Late Ordovician paleogeography of the South China blocks. (A) Location of South China on the reconstructed global paleogeographic map (after Scotese and McKerrow,1990). (B) Proposed
Late Ordovician paleogeography of China, after Mu et al. (1981). (C) Details of the Yangtze Platform, after Wang et al. (1997). Section localities: 1 — Wangjiawan; 2 — Nanbazi.
(Fig. 1). Regional palaeogeographical reorganization during the O–S
transition was coincided with eustatic sea-level changes (Zhang et al.,
2000; Chen et al., 2004). The Yangtze Platform was divided into the
western Upper Yangtze Platform and the eastern Lower Yangtze
Platform, which were separated by the Jiujiang Strait (Mu et al., 1981).
As sea level dropped, the Yangtze Platform, especially the Upper
Yangtze Platform, was largely enclosed (Wang et al., 1997; Chen et al.,
2004).
The graptolitic black shale facies predominated over the Yangtze
Platform during the O–S transition. Mu et al. (1981) first identified the
Ashgillian (Wufeng Formation) graptolite zones in this region. The
Hirnantian Substage defined here spans, in ascending order, from the
N. extraordinarius–N. ojsuensis zone to the N. persculptus zone (Rong
et al., 2000). Under this circumstance, the lowermost part of the
Longmaxi Formation that bears fossil elements of the N. persculptus
zone is placed in the uppermost Hirnantian Substage, above which is
the proposed boundary demarcating the Ordovician and Silurian
(Chen et al., 2000). In this paper, the biostratigraphic nomenclatures of
Chen et al. (2000) are adopted.
Biostratigratigraphically well-constrained sections from two different localities in the Yangtze basin were chosen for chemostratigraphic studies. They are the Wangjiawan section at Yichang, Hubei
province, and the Nanbazi section at Tongzi, Guizhou province. Rong
et al. (1999) and Chen et al. (1999) have suggested that the GSSP
candidate sections of the Hirnantian Substage can be established at
Wangjiawan, Yichang, Hubei province, and Honghuayuan at Tongzi,
Guizhou province. The better-preserved Nanbazi section that suffered
less weathering, ~ 500 m southeast of Honghuayuan, is thus selected
for the chemostratigraphic study in view of same depositional
successions. The depositional successions are dominated by black
shales intercalated with an argillaceous limestone horizon bearing
Hirantian fauna. Exact descriptions of these two sections (Wangjiawan
and Nanbazi) on their stratigraphy and palaeogeography can be seen
in Wang et al. (1993a,b), Chen et al. (1999), Rong et al. (1999) and Su
et al. (2003).
3. Analytical methods
Fresh samples were extracted and crushed into powders less than
200 mesh in size, ready for analysis of total organic carbon content
(TOC), δ13Corg and δ34Ssulfide ratios. TOC was measured by High TOC II
analyzer.
Sample splits (300 mg to 1.5 g) for δ13Corg analysis were firstly
dissolved with 5 N HCl in a centrifuge beaker to remove carbonates
through multiple acidification (at least two times) and subsequent
drying in the heating oven, and repeatedly rinsed with deionized
water to neutrality. The decalcified samples (30 to 110 mg) + CuO wire
(1 g) were added to a quartz tube, and combusted at 500 °C for 1 h and
850 °C for another 3 h. Isotopic ratios were analyzed using
cryogenically purified CO2 on a Finnigan MAT-253 mass spectrometer,
and reported in standand δ-notation relative to Vienna Peedee
Belmnite (VPDB) standard. Analytical precision for δ13Corg is better
than ± 0.06‰.
The sample split (1–2 g) for δ34Spy was treated with 100 ml CrCl2 +
10 ml alcohol in the vessel. The hydrogen sulfide was immediately
purged by the N2 stream, and was trapped and collected in the zinc
acetate (ZnAc) solution, in which 2% AgNO3 + 6 N NH4OH solutions were
added to precipitate Ag2S, then it was mixed with V2O5 + pure quartz
sands, and combusted up to 1050 °C in the presence of copper turnings
under vacuum to convert to SO2 gases. The purified SO2 was used to
analyze the sulfur isotopic ratio in the Finnigan Delta-S mass spectrometer. Sulfur isotopic ratios are expressed as standard δ-notation
34
D. Yan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 274 (2009) 32–39
Fig. 2. Carbon-sulfur isotopic chemostratigraphic profiles across the O–S boundary at Wangjiawan, Yichang, Hubei province, South China. Biostratigraphic framework is based on
Chen et al. (2000) and Su et al. (2003).
Fig. 3. Carbon-sulfur isotopic chemostratigraphic profiles across the O–S boundary at Nanbazi, Tongzi, Guizhou province, South China. Biostratigraphic framework is based on Chen et
al. (2000) and Su et al. (2003). See Fig. 2 for rock legends.
D. Yan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 274 (2009) 32–39
relative to Canon Diablo Troilite (CDT) standard. The analytical precision
is better than ±0.2‰.
4. Results
Forty-one and thirty-seven samples were analysed for TOC
abundance, δ13Corg and δ34Ssulfide values across the O–S boundary
succession at Wangjiawan and Nanbazi sections, respectively (Figs. 2
and 3; Tables 1 and 2).
At Wangjiawan, the TOC abundance varies generally between 0.35
and 9.32 wt.%, with the lowest abundance in the upper Hirnantian.
δ13Corg values vary commonly between −30.8 and −27.6‰ showing a
positive excursion starting from the lower P. pacicicus zone and
reaching the maximum (−27.6‰) in the uppermost part of the N.
extraordinarius zone in the upper Hirnantian at this locality. Going
upwards, the δ13Corg values decrease rapidly from the N. persculptus
zone in the uppermost Hirnantian through the P. acuminatus zone.
Further upwards, the δ13Corg values display relatively consistent
values between −30.1 and −30.0‰ from the O. vesiculosus through
the C. cyphus zones.
At Nanbazi, the TOC contents are highly variable, but are generally low in the Hirnantian horizon as well (Fig. 3; Table 2). The δ13Corg
values in this section show a positive excursion from −30.5‰ in the
D. complexus zone to the zenith of −26.6‰ near the top of the
N. extraordinarius zone of the upper Hirnantian (Fig. 3); this is fol-
35
lowed by a drastic drop in δ13Corg values of about 2.6‰ to the trough
at −29.2‰ in the uppermost Hirnantian, which bounce back slightly
to about −28‰ in the C. vesiculosuc zone and generally persist
upwards. Although a systematic offset of δ34Ssulfide value is apparent
from the two sections: relatively negative (avg. − 13.5‰) at
Wangjiawan and positive (avg. 1.3‰) at Nanbazi, positive δ34Ssulfide
excursions, starting from the P. pacificus zone and reaching the peak
in the N. extraordinarius zone (7.1‰ at Wangjiawan, 24.6‰ at
Nanbazi), are apparent both at these two sections (Figs. 2 and 3).
Going upwards, δ34Ssulfide values decrease sharply and reach the
trough (−27.5‰ at Wangjiawan, − 19.2‰ at Nanbazi) in the lower
Rhuddanian, which increase slightly upwards (Figs. 2 and 3). SEM
observations revealed that framboidal pyrite is the common and
major sulfur carrier in these rocks; these framboidal aggregates,
generally 2 to 5 μm in diameter, consist of cubic pyrite microcrystals
smaller than 1 μm (Fig. 4).
5. Discussion
Our data derived from two studied sections as described above
clearly show a positive δ13Corg excursion (“heavy-C event”), which is
followed by a rapid drop and a subsequent bounce across the O–S
boundary. A parallel variation in δ34Ssulfide values is well demonstrated
at the two sections as well (Figs. 2 and 3). The positive excursions in
δ13Corg and δ34Ssulfide values are estimated to have lasted for about
Table 1
Isotopic compositions of organic carbon and pyrite sulfur from Wangjiawan section, western Hubei, South China
Sample
Log height
No.
(m)
10
10+
11
12
13
14−
14
15
16
17
18
19
19+
20
21
22
23
24
25
26
27
27+
28
29
30
31
32
33
34
35
36
37
38
39
40
41
42
43
44
44+
45
0.05
0.2
0.35
0.6
0.85
1.05
1.15
1.4
1.55
1.7
1.85
2
2.08
2.15
2.3
2.4
2.55
2.65
2.75
2.8
2.85
2.9
2.93
3
3.02
3.07
3.15
3.25
3.35
3.45
3.6
3.8
4
4.2
4.4
4.65
4.95
5.2
5.5
5.65
5.8
Graptolite zone
Lithology
D. complexus
D. complexus
D. complexus
D. complexus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
N. extraordinarius–N. ojsuensis
N. extraordinarius–N. ojsuensis
N. extraordinarius–N. ojsuensis
N. extraordinarius–N. ojsuensis
N. extraordinarius–N. ojsuensis
N. extraordinarius−N. ojsuensis
N. extraordinarius–N. ojsuensis
N. extraordinarius−N. ojsuensis
N. extraordinarius–N. ojsuensis
N. extraordinarius–N. ojsuensis
N. persculptus
N. persculptus
N. persculptus
N. persculptus
A. ascensus
A. ascensus
P. acuminatus
P. acuminatus
C. vesiculosus
C. vesiculosus
C. vesiculosus
C. cyphus
C. cyphus
C. cyphus
C. cyphus
C. cyphus
C. cyphus
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Carbonaceous mudstone
Carbonaceous mudstone
Carbonaceous mudstone
Carbonaceous mudstone
Carbonaceous mudstone
Carbonaceous mudstone
Argillaceous limestone
Argillaceous limestone
Argillaceous limestone
Argillaceous limestone
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Mudstone
Mudstone
Mudstone
Silty mudstone
Silty mudstone
Silty mudstone
δ13Corg
δ34Ssulfide
(%)
(‰, VPDB)
(‰, CDT)
2.4
2.45
2.58
4.48
4.2
4.19
4.2
3.05
3.64
4.73
5.31
2.8
7.18
9.32
7.2
4.81
4.15
3.73
2.93
1.74
0.35
0.6
0.62
0.55
6.65
7.72
7.65
7.69
4.98
5.76
4.71
2.86
2.68
3.97
3.45
4.05
4.98
4.57
4.98
5.07
5.18
−30.74
−30.79
−30.83
−30.61
−30.69
−30.28
−29.88
−30.45
−29.62
−30.72
−30.31
−29.31
−29.44
−30.01
−29.64
−29.15
−29.22
−29.4
−29.33
−29.09
−28.68
− 27.72
− 27.64
−29.21
−28.94
−29.04
−30.17
−30.01
−30.29
−30.33
−30.48
−28.96
−30.13
−30.16
−29.45
−30.13
−30.07
−30.01
−30.06
−29.98
−30.01
− 16.36
−17.28
−18.72
−17.7
− 22.34
− 22.51
− 23.61
−18.7
− 19.16
−18.73
−17.15
−11.8
−8.01
−7.33
−1.18
−9.52
−1.93
− 19.35
7.11
− 16.06
1.48
6.22
−1.74
5.34
− 14.1
−18.87
−21.88
−17.67
−17.02
− 16.83
− 19.12
−27.46
− 22.32
−21.18
−21.85
− 15.92
−17.96
− 13.06
−4.75
−7.15
−9.06
TOC
36
D. Yan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 274 (2009) 32–39
Table 2
Isotopic compositions of organic carbon and pyrite sulfur from Nanbazi section, northern Guizhou, South China
Sample
Log height
No.
(m)
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27−
27
28
29
30
31
32
33
34
35
36
37
38
39
40
41
0.3
0.6
1.15
1.95
2.85
3.75
4.55
5.25
5.75
6.25
6.55
6.8
6.95
7.15
7.5
8.05
8.45
9
9.5
10.05
10.25
10.55
10.8
11.1
11.6
11.9
12.2
13.25
13.5
13.85
14.25
14.45
14.95
15.45
16.45
17.6
18.6
Graptolite zone
Lithology
D. complexus
D. complexus
D. complexus
D. complexus
D. complexus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
P. pacificus
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
N. extraordinarius–N.
?
?
?
C. vesiculosus
C. vesiculosus
C. vesiculosus
C. cyphus
C. cyphus
C. cyphus
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Shale
Argillaceous limestone
Argillaceous limestone
Argillaceous limestone
Argillaceous limestone
Argillaceous limestone
Mudstone
Mudstone
Argillaceous limestone
Argillaceous limestone
Mudstone
Mudstone
Mudstone
Argillaceous limestone
Argillaceous limestone
Argillaceous limestone
Argillaceous limestone
Mudstone
Mudstone
Mudstone
Silty mudstone
Silty mudstone
Silty mudstone
Silty mudstone
Silty mudstone
Silty mudstone
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
ojsuensis
1 m.y., given that each graptolite zone lasted for 0.5 m.y. in duration
(Mu et al., 1981; Brenchley et al., 1994).
The positive δ13Corg excursion and subsequent variations across
the O–S boundary reported here are unlikely of local signatures, since
this positive excursion both in δ13Ccarb and δ13Corg values associated
with the Hirnantian interval has been widely reported from a number
of sections around the world, including Siljan area, central Sweden
(Marshall and Middleton, 1990), Mackenzie Mountains (Wang et al.,
1993a,b) and Anticosti Island, Canada (Long, 1993), Scotland (Underwood et al., 1997), South China (Wang et al., 1997), Nevada, USA
(Finney et al., 1999; Kump et al., 1999), and Estonia and Latvia
(Brenchley et al., 2003; Kaljo et al., 2004). Although this δ13C excursion
is widespread, the timing of the peak may not be completely
synchronous between different regions (Melchin and Holmden,
2006). It is also interesting to note that a δ13Corg excursion of similar
to somewhat larger magnitude than the δ13Ccarb excursion is present
(Melchin and Holmden, 2006).
The abrupt shift in carbon isotope values indicates a significant
change in the relative carbon fluxes between ocean water, organic
matter and sediments. Any model of carbon cycle must be consistent
with the temporal relationships among carbon cycle, sea-level fall,
and temperature change identified here (Brenchley et al., 2003). Kump
et al. (1999) suggested that the positive δ13C excursion was the result
of increased carbonate-platform weathering during glacioeustatic
sea-level lowstand. Whether or not the changes in weathering
processes could trigger the glaciation event, however, is still highly
controversial (Melchin and Holmden, 2006). In an alternative model,
Brenchley et al. (1994, 2003) proposed that the positive δ13C shift was
caused by enhanced burial of carbon or enhanced productivity leading
TOC
δ13Corg
δ34Ssulfide
(wt.%)
(‰, VPDB)
(‰, CDT)
5.72
4.74
5.06
4.95
4.21
4.74
5.25
5.25
6.63
10.42
10.21
6.62
4.21
1.98
3.02
1.23
1.34
1.57
2.09
0.38
0.76
1.76
0.51
0.56
0.26
0.14
0.12
0.09
8.45
2.14
1.52
0.47
0.6
0.51
0.37
1.31
0.51
−30.48
−30.46
−30.42
−30.29
−30.26
−30.29
−29.94
−29.87
−29.74
−29.12
−28.9
−28.8
−28.69
−28.7
−28.82
−28.58
−28.47
−28.39
−28.5
−28.08
−28.09
−28
−27.1
−28.08
−26.85
−27.17
−26.63
−27.48
−28.99
−29.23
−29.05
−28.25
−28.34
−28.03
−28.02
−28.29
−28.17
− 13.52
− 6.29
− 11.21
–
–
− 15.83
− 13.02
–
− 11.01
0.12
0.93
8.98
3.21
5.98
16.58
14.78
17.6
11.01
20.79
10.37
9.37
18.33
20.45
24.59
14.21
13.21
9.24
4.66
− 15.79
− 18.7
− 19.21
− 15.32
–
− 13.17
− 5.14
− 11.37
− 12.21
to preferential removal of 12C at least from surface waters. The low
abundance of organic carbon within this excursion horizon (Figs. 2
and 3) excludes the possibility of high productivity during the
deposition. On the other hand, the high abundance of organic carbon
below the Hirnantian horizon points to an enhanced burial of organic
matter likely in much more anoxic water columns (Yan et al., 2008),
during which the warm climate (high pCO2 level) could have
accelerated photosynthetic carbon fixation of marine phytoplankton,
leading to lower values of δ13Corg (Freeman and Hayes, 1992; Rau et al.,
1997) as observed (Figs. 2 and 3). The progressive increase in burial of
organic matter could have led to substantial removal of 12C from the
carbon reservoir, resulting in the climate cooling (even glaciation) and
a positive excursion of δ13C (Kump and Arthur, 1999) during the
Hirnantian. During the Hirnantian glaciation, the low pCO2 level, on
the other hand, could have decreased photosynthetic carbon fixation
of marine phytoplankton, through which less 12C would be incorporated into the organic carbon, leading to the heavy δ13Corg values
(Freeman and Hayes, 1992; Rau et al., 1997). The subsequent rapid rise
in sea-level, and recurrences of oceanic anoxia and warm climate,
probably triggered by unidentified forcing, resumed the scenarios as
those prior to Hirnantian glaciation, in which the organic matter was
massively accumulated and buried in parallel with a sharp negative
excursion of δ13Corg as observed (Figs. 2 and 3) as a result of the
acceleration of carbon fixation of marine phytoplankton under a high
pCO2 level.
The δ34Ssulfide values of sulfides precipitated on the earth surface
are dependent on the fractionation of bacterial sulphate reduction in
water media, which is commonly related to the oxygen level, sulphate
and iron concentrations of water columns (Canfield and Thamdrup,
D. Yan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 274 (2009) 32–39
37
Fig. 4. Scanning electron microscope (SEM) photomicrographs of pyrite from studied samples: (A) cluster of well developed octahedral crystals in sample 5 from Wangjiawan;
(B) framboidal pyrite made up of uniform microcrystals in sample 23 from Wangjiawan; (C) pyrite framboids in sample 38 from Wangjiawan; (D) pyrite framboids in sample 18 from
Nanbazi.
1994; Habicht et al., 1998; Strauss, 1999). The offset of δ34Ssulfide values
between Wangjiawan and Nanbazi (Figs. 2 and 3) may reflect the
differences in depositional setting. Paleogeographically, the Yangtze
Platform was a sort of epeiric shelf during the O–S transition, on which
numerous sub-depressions and sub-highs were present (Mu et al.,
1981; Chen et al., 2004). The Nanbazi was located in a sub-depression
in the innermost part of the shelf, while Wangjiawan was located on
the outer part of the shelf facing the northern oceanic basin (Fig. 1);
therefore, the Nanbazi was more restricted in water circulation, in
which the progressive bacterial sulphate reduction could have
resulted in the depletion of sulfate in the water column and increased
influences of sulphate sulfur isotopes upon the sulfide sulfur isotope
compositions, thereby leading to heavier δ34Ssulfide values (Canfield
and Thamdrup, 1994; Habicht and Canfield, 1997) as demonstrated in
this study. In addition, a relatively high oxygen level in water columns
at Nanbazi due to its proximity to the source area and shallower water,
which could have depressed the bacterial sulphate reduction as well,
may also contribute, more or less, to the heavy δ34Ssulfide values there.
The similar patterns in δ34Ssulfide changes: a positive excursion
followed by a rapid negative shift during the O–S transition, both at
Nanbazi and Wangjiawan (Figs. 2 and 3) could be a reflection of similar
oceanic environmental changes, likely in response to changes in
oxygen level in water columns that were further controlled by eustatic
sea-level fluctuations and climate changes. Studies on C–S–Fe relationship of organic-rich sediments spanning O–S transition revealed a
stratified anoxic water column on the shelf during the mid-Ashgill
interval to an oxygenated water condition during the mid-early
Hirnantian, which was followed by a rapid recurrence of anoxic
water column during the latest Hirnantian–earliest Rhuddanian
interval (Yan et al., 2008). The temporal coincidence of light δ34Ssulfide
pattern with the anoxic periods, and heavy δ34Ssulfide pattern with
ventilated periods, suggests that the oxygen level of bottom waters
could have played a major role in controlling the fractionation of
bacterial sulphate reduction in water–sediment interface. During
the oceanic anoxia, the bacterial sulphate reduction was greatly
enhanced, through which 32S was increasingly incorporated into the
sulfides precipitating from bottom waters, leading to the light
δ34Ssulfide values. On the contrary, during the ventilated period, i.e.,
during the mid-early Hirnantian interval, the bacterial sulphate
reduction could have substantially decreased in the bottom water,
leading to the heavy δ34Ssulfide values due to less incorporation of
32
S into the sulfides (Canfield and Thamdrup, 1994; Habicht and
Canfield, 1996).
6. Implications for the Late Ordovician extinction
The latest Ordovician biotic crisis was marked by two-phase
extinction events (Brenchley et al., 1995). The first strike occurred
during the glacial period (or a eustatic sea-level fall) coincident with
the positive shift in δ13C values of carbonates and organic carbon; the
second wave of extinctions occurred during the subsequent rise in
sea-level marked by a large negative shift in δ13C values (Marshall
et al., 1997). Our δ13Corg and δ34Ssulfide isotopic data also provide a
good constraint on the timing of the two-phase extinctions and
relevant oceanic and climatic changes. The rapid shift from warm to
cold climate accompanied with eustatic sea-level fall in the beginning
of Hirnantian age could have exerted severe ecological stresses on
those organisms, both benthic and planktonic dwellers, who had been
adapting to tropical warm waters, leading to their massive demise.
The subsequent sharp rise in sea-level and return of warm climate
38
D. Yan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 274 (2009) 32–39
accompanied with oceanic anoxia gave another detrimental strike for
those organisms, especially the benthic habitats, i.e., the Hirnantian
fauna, that had been adapting to cool temperature and ventilated
shallow-water conditions, resulting in the second wave of severe
declines of these particular fauna. In general, the organisms can
gradually get accustomed to slow long-term climatic changes, but
they cannot readily adjust their physiology to frequent short-term
climatic changes, which could inhibit the origination of new species
and their radiation (Clarke, 1993). Moreover, the oceanic anoxia is also
harmful for the biotic colonization mainly because of the presence of
abundant toxic H2S in the water columns (Kump et al., 2005; Algeo
et al., 2007). In view of Fe depletion in the anoxic water bodies during
the O–S transition (Yan et al., 2008), significant amounts of H2S gases
produced through bacterial sulphate reduction during the oxidation of
organic matter could not be fully sequestrated to form sulfides (mainly
pyrite), so that they would be released into the photic zone and even
atmosphere, resulting in the photic euxinia (Kump et al., 2005); this
scenario would be much more harmful, both for the benthic and
planktonic fauna. Therefore, the frequent large-scale climate fluctuations, together with multiple oceanic anoxia (particularly photic
euxinia) and sea-level fluctuations, may account for the stepwise
massive demises during the O–S transition.
Acknowledgements
We thank Fusong Zhang and Lianjun Feng (IGGCAS) for access to
the stable isotope laboratory and assistance in stable isotope analysis.
This research was supported by the National Key Basic Research
Project (973 Project) through grant 2005CB422101 and NSFC-SINOPEC
United Foundation through grant 40839907. Thanks to Ben Gill,
Charlie Underwood, Junxuan Fan and Finn Surlyk for their constructive comments to improve the original manuscript.
References
Algeo, T.J., Ellwood, B., Nguyen, T.K.T., Rowe, H., Maynard, J.B., 2007. The Permian–
Triassic boundary at Nhi Tao, Vietnam: evidence for recurrent influx of sulfidic
water masses to a shallow-marine carbonate platform. Palaeogeogr. Palaeoclimat.
Palaeoecol. 252, 304–327.
Brenchley, P.J., 1988. Environmental changes close to the Ordovician–Silurian boundary.
Geology 43, 377–385.
Brenchley, P.J., Carden, G.A.F., Marshall, J.D., 1995. Environmental changes associated
with the “first strike” of the Late Ordovician mass extinction. Mod. Geol. 20, 69–82.
Brenchley, P.J., Carden, G.A.F., Hint, L., Kaljo, D., Marshall, J.D., Martma, T., Meidla, T.,
Nolvak, J., 2003. High-resolution stable isotope stratigraphy of Upper Ordovician
sequences: constraints on the timing of bioevents and environmental changes
associated with mass extinction and glaciation. Bull. Geol. Soci. Am. 115, 89–104.
Brenchley, P.J., Marshall, J.D., Carden, G.A.F., Robertson, D.B.R., Long, D.G.F., Meidla, T.,
Hints, L., Anderson, T.F., 1994. Bathymetric and isotopic evidence for a short-lived
Late Ordovician glaciation in a greenhouse period. Geology 22, 295–298.
Burdett, J.W., Arthur, M.A., Richardson, M., 1989. A Neogene seawater sulfur isotope age
curve from calcareous pelagic microfossils. Earth Planet. Sci. Lett. 94, 189–198.
Canfield, D.E., Thamdrup, B., 1994. The production of 34S depleted sulfide during
bacterial disproportionation of elemental sulfur. Science 266, 1973–1975.
Chen, X., Rong, J., Mitchell, C.E., Harper, D.A.T., Fan, J., Zhang, Y., Zhan, R., Wang, Z., Wang,
Z., Wang, Y., 1999. Stratigraphy of the Hirnantian Substage from Wangjiawan,
Yichang, W. Hubei and Honghuayuan, Tongzi, N. Guizhou, China. In: Kraft, P., Fatka,
O. (Eds.), Quo vadis Ordovician?, vol. 43 1/2. Acta Universitatis Carolinae–Geologica,
pp. 233–236.
Chen, X., Rong, J., Fan, J., Zhan, R., Zhang, Y., Li, R., Wang, Y., Mitchell, C.E., Harper, D.A.T.,
2000. A global correlation of biozones across the Ordovician–Silurian boundary.
Acta Palaeont. Sin. 39, 100–114.
Chen, X., Rong, J., Li, Y., Boucot, A.J., 2004. Facies patterns and geography of the Yangtze
region, South China, through the Ordovician and Silurian transition. Palaeogeogr.
Palaeoclimat. Palaeoecol. 204, 353–372.
Clarke, A., 1993. Temperature and extinction in the sea: a physiologist's view.
Palaeobiology 19, 499–518.
Fike, D.A., 2007. Oxidation of the Ediacaran Ocean. Nature 444, 744–747.
Finney, S.C., Berry, W.B.N., Cooper, J.D., Sweet, W.C., Jacobson, S.R., Soufiane, A., Achab,
A., Npble, P.J., 1999. Late Ordovician mass extinction: a new perspective from
stratigraphic sections in central Nevada. Geology 27, 215–218.
Freeman, K.H., Hayes, J.M., 1992. Fractionation of carbon isotopes by phytoplankton and
estimates of ancient pCO2 levels. Glob. Biogeochem. Cycles 6, 629–644.
Geldsetzer, H.H.J., Goodfellow, W.D., Mclaren, D.J., Orchard, M.J., 1987. Sulfur-isotope
anomaly associated with the Frasnian–Famennian extinction, Medicine Lake,
Alberta, Canada. Geology 15, 393–396.
Gill, B.C., Lyons, T.W., Saltzman, M.R., 2007. Parallel, high-resolution carbon and sulfur
isotope records of the evolving Paleozoic marine sulfur reservoir. Palaeogeogr.
Palaeoclimatol. Palaeoecol. 256, 156–173.
Goodfellow, W.D., Nowlan, G.S., McCracken, A.D., Lenz, A.C., Gregoire, D.C., 1992.
Geochemical anomalies near the Ordovician–Silurian boundary, Northern Yukon
Territory, Canada. Hist. Biol. 6, 1–23.
Habicht, K.S., Canfield, D.E., 1996. Sulphur isotope fractionation in modern microbial
mats and the evolution of the sulphur cycle. Nature 382, 342–343.
Habicht, K.S., Canfield, D.E., 1997. Sulfur isotope fractionation during bacterial sulfate
reduction in organic-rich sediments. Geochim. Cosmochim. Acta 61 (24),
5351–5361.
Habicht, K.S., Canfield, D.E., Rethmeier, J., 1998. Sulfur isotope fractionation during
bacterial reduction and disproportionation of thiosulfate and sulfite. Geochim.
Cosmochim. Acta 62 (15), 2585–2595.
Joachimski, M.M., Henning, C.O., Pancost, R.D., Strauss, H., Freeman, K.H., Littke, R.,
Damste, J.S.S., Racki, G., 2001. Water column anoxia, enhanced productivity and
concomitant changes in δ13C and δ34S across the Frasnian–Famennian boundary.
Chem. Geol. 175, 109–131.
Kajiwara, Y., Yamakita, S., Ishida, K., Ishiga, H., Imai, A., 1994. Development of a largely
anoxic stratified ocean and its temporary massive mixing at the Permian/Triassic
boundary supported by the sulfur isotopic record. Palaeogeogr. Palaeoclimat.
Palaeoecol. 111, 367–379.
Kaljo, D., Hints, L., Martma, T., Nõlvak, J., Oraspõld, A., 2004. Late Ordovician carbon
isotope trend in Estonia, its significance in stratigraphy and environmental analysis.
Palaeogeogr. Palaeoclimat. Palaeoecol. 210, 165–185.
Kump, L.R., Arthur, M.A., 1999. Interpreting carbon-isotope excursions: carbonates and
organic matter. Chem. Geol. 161, 181–198.
Kump, L.R., Arthur, M.A., Patzkowsky, M.E., Gibbs, M.T., Pinkus, D.S., Sheehan, P.M., 1999.
A weathering hypothesis for glaciation at high atmospheric pco2 during the Late
Ordocician. Palaeogeogr. Palaeoclimat. Palaeoecol. 152, 173–187.
Kump, L.R., Pavlov, A., Arthur, M.A., 2005. Massive release of hydrogen sulphide to the
surface ocean and atmosphere during intervals of oceanic anoxia. Geology 33,
397–400.
Long, D.G.F., 1993. Oxygen and carbon isotopes and event stratigraphy neat the
Ordovician–Silurian boundary, Anticosti Island, Quebec. Palaeogeogr. Palaeoclimat.
Palaeoecol. 104, 49–59.
Marshall, J.D., Middleton, P.D., 1990. Changes in marine isotopic composition and the
Late Ordovician glaciation. J. Geol. Soc., Lond. 147, 1–4.
Marshall, J.D., Brenchley, P.J., Mason, P., Wolff, G.A., Astini, R.A., Hints, L., Meidla, T., 1997.
Global carbon isotopicevents associated with mass extinction and glaciation in the
late Ordovician. Palaeogeogr. Palaeoclimat. Palaeoecol. 132, 195–210.
Melchin, M.J., Holmden, C., 2006. Carbon isotope chemostratigraphy in Arctic Canada:
sea-level forcing of carbonate platform weathering and implications for Hirnantian
global correlation. Palaeogeogr. Palaeoclimat. Palaeoecol. 234, 186–200.
Metcalfe, I., 1994. Late Palaeozoic and Mesozoic palaeogeography of eastern Pangaea
and Tethys. Can. Soc. Pet. Geol., Mem. 17, 97–111.
Mu, E., Li, J., Ge, M., Chen, X., Ni, Y., Lin, Y., 1981. Late Ordovician paleogeography of south
China. Acta Stratigr. Sin. 5, 165–170.
Newton, R.J., Pevitt, E.L., Wignall, P.B., Bottrell, S.H., 2004. Large shifts in the isotopic
composition of seawater sulphate across the Permo-Triassic boundary in northern
Italy. Earth Planet. Sci. Lett. 218, 331–345.
Rau, G.H., Riebesell, U., Wolf-Gladrow, D., 1997. CO2,aq-dependent photosynthetic 13C
fractionation in the ocean: a model versus measurements. Glob. Biogeochem.
Cycles 11, 267–278.
Riccardi, A.L., Kump, L.R., Arthur, M.A., 2007. Carbon isotopic evidence for chemocline
upward excursions during the end-Permian event. Palaeogeogr. Palaeoclimatol.
Paleoecol. 248, 73–81.
Rong, J., Chen, Xu., Harper, D.A.T., Mitchell, C.E., 1999. Proposal of a GSSP candidate
section in the Yangtze Platform region, S. China, for a new Hirnantian boundary
stratotype. In: Kraft, P., Fatka, O. (Eds.), Quo vadis Ordovician?, vol. 43 1/2. Acta
Universitatis Carolinae–Geologica, pp. 77–80.
Rong, J., Chen, X., Harper, D.A.T., Mitchell, C.E., 2000. Proposal of a GSSP candidate
section in South China for a new Hirnantian boundary stratotype. J. Stratigr. 24,
176–181.
Scotese, C.R., McKerrow, W.S., 1999. Revised world maps and introduction. In:
McKerrow, W.S., Scotese, C.R. (Eds.), Palaeozoic Palaeogeography and Biogeography,
vol. 12. Geol. Soc. London, Mem. pp. 1–21.
Strauss, H., 1997. The isotopic composition of sedimentary sulfur through time.
Palaeogeogr. Palaeoclimat. Palaeoecol. 132, 97–118.
Strauss, H., 1999. Geological evolution from isotope proxy signals — sulfur. Chem. Geol.
161, 89–101.
Su, W.B., He, L.Q., Wang, Y.B., Gong, S.Y., Zhou, H.Y., 2003. K-Bentonite beds and highresolution integrated stratigraphy of the uppermost Ordovician Wufeng and the
lowest Silurian Longmaxi formations in South China. Sci. China 46 (11), 1121–1133.
Underwood, C.J., Crowley, S.F., Marshall, J.D., Brenchley, P.J., 1997. High-resolution
carbon isotope stratigraphy of the basal Silurian stratotype (Dob's Linn, Scotland)
and its global correlation. J. Geol. Soc. Lond. 154, 709–718.
Wang, K., Chatterton, B.D.E., Attrep, M., Orth, C.J., 1993a. Late Ordovician mass extinction
in the Selwyn basin, northwestern Canada: geochemical, sedimentological and
paleontological evidence. Can. J. Earth Sci. 30, 1870–1880.
Wang, K., Orth, C.J., Attrep, J.M., Chatterton, B.D.E., Wang, X., Li, J., 1993b. The great latest
Ordovician extinction on the South China Plate: chemostratigraphic studies of the
Ordovician–Silurian boundary interval on the Yangtze Platform. Palaeogeogr.
Palaeoclimat. Palaeoecol. 104, 61–79.
Wang, K., Chatterton, B.D.E., Wang, Y., 1997. An organic carbon isotope record of late
Ordovician to early Silurian marine sedimentary rocks, Yangtze sea, South China:
D. Yan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 274 (2009) 32–39
implications for CO2 changes during the Hirnantian glaciation. Palaeogeogr.
Palaeoclimat. Palaeoecol. 132, 147–158.
Wilde, P., Berry, W.B.N., Quinby-Hunt, M.S., Orth, C.J., Quintana, L.R., Gilmore, J.S., 1986.
Iridium abundances across the Ordovician–Silurian stratotype. Science 233,
339–341.
39
Yan, D., Chen, D., Wang, Q., Wang, J., Chu, Y., 2008. Environment redox changes of the Yangtze
Sea during the Ordo-Silurian transition. Acta Geol. Sin.-English Eedition 82 (3), 679–689.
Zhang, T.S., Kershaw, S., Wan, Y., Lan, G.Z., 2000. Geochemical and facies evidence for
palaeoenvironmental change during the Late Ordovician Hirnantian glaciation in
South Sichuan Province, China. Glob. Planet. Change 24, 133–152.