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Journal of Earth Science, Vol. 26, No. 2, p. 166–180, April 2015
Printed in China
DOI: 10.1007/s12583-015-0523-1
ISSN 1674-487X
Rapid Carbonate Depositional Changes Following the
Permian-Triassic Mass Extinction: Sedimentary
Evidence from South China
Li Tian1, Jinnan Tong*1, David Bottjer2, Daoliang Chu1, Lei Liang1, Huyue Song1, Haijun Song1
1. State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Wuhan 430074, China
2. Department of Earth Sciences, University of Southern California, Los Angeles CA90089, USA
ABSTRACT: Various environmental changes were associated with the Permian-Triassic mass extinction at 252.2 Ma. Diverse unusual sediments and depositional phenomena have been uncovered as responses to environmental and biotic changes. Lithological and detailed conodont biostratigraphic correlations within six Permian-Triassic boundary sections in South China indicate rapid fluctuations in
carbonate deposition. Four distinct depositional phases can be recognized: (1) normal carbonate deposition on the platform and slope during the latest Permian; (2) reduced carbonate deposition at the onset of the main extinction horizon; (3) expanded areas of carbonate deposition during the Hindeodus
changxingsensis Zone to the H. parvus Zone; and (4) persistent mud-enriched carbonate deposition in
the aftermath of the Permian-Triassic transition. Although availability of skeletal carbonate was significantly reduced during the mass extinction, the increase in carbonate deposition did not behave the
same way. The rapid carbonate depositional changes, presented in this study, suggest that diverse environmental changes played key roles in the carbonate deposition of the Permian-Triassic mass extinction
and onset of its aftermath. An overview of hypotheses to explain these changes implies enhanced terrestrial input, abnormal ocean circulation and various geobiological processes contributed to carbonate
saturation fluctuations, as the sedimentary response to large volcanic eruptions.
KEY WORDS: Permian-Triassic, mass extinction, carbonate, sedimentary response, environmental change.
0
INTRODUCTION
The most distinct conversion of biotic environments in the
Phanerozoic happened during the Paleozoic-Mesozoic transition, associated with the Permian-Triassic (P-T) mass extinction (Peters, 2008; Erwin, 1994). Lethally hot temperatures
caused by high pCO2 as well as H2S in the anoxic ocean are
thought to be responsible for extinction of 96% marine invertebrates (Piestch and Bottjer, 2014; Song H J et al., 2014, 2012a;
Tian et al., 2014a; Sun Y D et al., 2012; Bottjer et al., 2008;
Knoll et al., 2007; Grice et al., 2005; Raup, 1979). The eruption
of the Siberian Traps released over 105 Gt CO2 and other gases
into the atmosphere and triggered many other devastating environmental events, such as rapid temperature changes (volcanic winter to the hothouse), acid rain and ocean acidification
as well as atmospheric O2 depletion and ocean anoxia/dysoxia
(Svensen et al., 2009; Wignall, 2001; Campbell et al., 1992).
These environmental changes forced the vertebrates on land to
migrate to high latitude areas and suppressed marine metazoans
to habitation in minimal refuges (Piestch and Bottjer, 2014;
Song H J et al., 2014; Sun Y D et al., 2012).
*Corresponding author: [email protected]
© China University of Geosciences and Springer-Verlag Berlin
Heidelberg 2015
Manuscript received Septermber 12, 2014.
Manuscript accepted Februry 15, 2015.
Associated with the mass extinction, oceanic depositional
systems returned to conditions similar to Precambrian-Cambrian
oceans, indicated by the global occurrence of unusual sediments, such as microbialites, stromatolites, giant ooids, sea
floor fan precipitates and wrinkle structures (Li et al., 2013;
Baud et al., 2007; Knoll et al., 2007; Pruss et al., 2006). This
turnover has been interpreted as the sedimentary response to
ecosystem changes, attributed to unusual seawater chemistry,
perturbation of bioturbation and grazing pressures, and other
changes in actualistic sedimentologic processes (Woods, 2014;
Pruss et al., 2006).
For example, end-Permian metazoan reef or reef like bioclastic limestones were replaced by basal Triassic microbialites
(Yang et al., 2011; Xie et al., 2010; Wang et al., 2009; Payne et
al., 2007), coinciding with the mass extinction. Kershaw et al.
(2007) attributed the growth of basal Triassic microbialites to
upwelling. These unusual sediments also were thought to be
caused by the overturn of carbonate saturation following ocean
acidification (Payne et al., 2010; Kump et al., 2009). Along with
the distinct sedimentary turnover, biotic changes also occurred,
leading to a significant decrease in skeletons through the P-T
transition (Tian et al., 2014b; Wang et al., 2009; Kaiho et al.,
2006). Knoll et al. (2007, 1996) proposed that photosynthesis and
sulfate reduction contributed to the extinction of calcified metazoan and formation of unusual sediments in the overturned
oceans. The carbonate deposition process could be similar with
the modern black sea (Woods et al., 1999; Kempe, 1990).
Tian, L., Tong, J. N., Bottjer, D., et al., 2015. Rapid Carbonate Depositional Changes Following the Permian-Triassic Mass
Extinction: Sedimentary Evidence from South China. Journal of Earth Science, 26(2): 166–180. doi:10.1007/s12583-015-0523-1
Rapid Carbonate Depositional Changes Following the Permian-Triassic Mass Extinction
Although all these hypotheses can explain the sedimentary and
biotic changes, an overview of the observed evidence and fundamental depositional processes needs further investigation.
Here, by reexamining the sedimentary changes of six
Permian-Triassic boundary (PTB) sections, which were located
from onshore to offshore during the P-T transition in South
China (Figs. 1, 2), rapid fluctuations of the carbonate factory
during the P-T transition are demonstrated. In addition a critical
review of the ocean acidification and upwelling theories also
provides the basis for proposing a new marine oxygenation
theory for precipitation of these unusual carbonate sediments
and fluctuations of the carbonate deposition during this time.
1
GEOLOGICAL SETTINGS AND MATERIALS
The Zhongzhai Section is located in the northeast part of
Zhongzhai, Liuzhi County, southwestern Guizhou Province.
The studied PTB of the Zhongzhai Section is comprised of the
upper part of the Longtan Formation and the basal Yelang Formation. The upper part of the Longtan Formation is composed
of calcareous mudstone and sandstone while the basal Yelang
Formation is composed of mudstone and siltstone. There are
two limestone beds intercalated between the top of the Longtan
Formation and the basal Yelang Formation (Fig. 2). The basal
4 cm of the boundary limestone is abundant in diverse metazoan fossils, including fusulinid fragments, foraminifers,
brachipoids, and molluscs (Figs. 3a, 3b, 3c). The upper part of
the boundary limestone is composed of abundant quartz grains
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and a few small foraminifers. Examination of polished slabs of
PTB limestone reveals that the grains in the basal 1 cm are
significantly larger than the upper part (Fig. 4a). No obvious
laminated layers or bioturbation have been observed.
The Dajiang Section is located on the Great Bank of
Guizhou, which was an isolated platform in the Nanpanjiang
Basin during the Permian to Triassic (Fig. 1). The Lower Triassic Daye Formation truncates the Upper Permian Wujiaping
Formation with an erosional surface (Fig. 2a). The Wujiaping
Formation is dominated by diverse metazoans, including algae,
foraminifers, brachiopods, and molluscs (Figs. 5a, 5b). The
base of the Daye Formation is composed of 17 m of thick massive microbialites, from which calcified cyanobacteria can be
observed in thin-sections (Fig. 5c). There are also several shelly
beds, dominated by ostracods (Fig. 5d), preserved within the
microbialites.
The Yangou Section is situated 30 km east of Jingdezhen
City, Jiangxi Province. The Upper Permian Changxing Formation underlies the basal Lower Triassic Daye Formation (Fig. 2)
in a continuous carbonate P-T succession (Song H J et al., 2012b;
Sun D Y et al., 2012). The studied upper part of the Changxing
Formation is composed of thick-massive bioclastic grainstone,
abundant in fusulinids, other foraminifers, algae, brachiopods
and molluscs (Fig. 6a). In contrast, skeletons are poor in the basal
Daye Formation thin limestone beds except for a few ostracod
and gastropod-enriched thin muddy-limestone beds (Figs. 6c, 6d).
There are also 2 distinct oolitic limestone beds observed
Figure 1. Paleogeographic map of South China during the P-T transition (after Yin et al., 2014). The studied sections (marking by red dots) were located in different facies. ZZ. Zhongzai Section; DJ. Dajiang Section; YG. Yangou Section; MS. Meishan Section; XK. Xiakou Section; CH. Chaohu Section; GL Bay. Guangyuan-Liangping Bay; XL Bay. Xiakou-Lichuan Bay;
NMBY. north marginal basin of Yangtze Platform; HGG Basin. Hunan-Guizhou-Guangxi Basin; ZFG clastic region; Zhejiang-Fujian-Guangdong clastic region.
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Li Tian, Jinnan Tong, David Bottjer, Daoliang Chu, Lei Liang, Huyue Song and Haijun Song
Figure 2. The correlation of the studied PTB sections and photos of the outcrops. (a) Photographs showing the erosion surface
between the latest Permian bioclastic limestone with the basal Triassic microbialites at the Dajiang Section; (b) the P-T
boundary limestone at the Zhongzhai Section; (c) the outcrops of the Yangou Section; (d) the P-T boundary at the Meishan
sections; (e) the Chaohu Section; (f) the P-Tr boundary at the Xiakou Section. PTBCL. Permian-Triassic boundary limestone.
at the base of the Daye Formation (Tian et al., 2014b). The
ooids were recrystallized heavily, leaving only the cortex preserved, cemented by spar (Fig. 6b).
The GSSP at the Meishan Section is located in Changxing
County, Zhejiang Province. The medium-thick bioclastic limestone dominating the Changxing Formation is overlain by the
muddy limestone and mudstone which dominates the Yinkeng
Formation (Fig. 2). At the top of the Changxing Formation
there are abundant foraminifers, brachiopods and molluscs
(Figs. 7a, 7b) while there are a few foraminifers preserved in
the basal Yinkeng Formation (Fig. 7c). The PTB limestone
(Bed 27) is a 16 cm thick limestone, sandwiched by clays and
mudstone (Fig. 2d). The PTB is located at the base of the Bed
27c, indicated by the first occurrences of Hindeodus parvus
(Jiang et al., 2007; Yin et al., 2001). By checking the polished
PTB limestone, bioturbation in 27b is remarkably different with
the laminated layers in 27c (Fig. 4b).
The Xiakou Section is 40 km northwest of Yichang City,
Hubei Province. It was situated at the northern margin of the
Yangtze platform during the end-Permian to Middle Triassic.
The PTB succession consists of the uppermost Permian Dalong
Formation, dominated by siliceous mudstone/limestone, and
the Lower Triassic Daye Formation, dominated by
thin-medium limestone, with interbedded claystone (Fig. 2).
The PTB is thought to be located in the middle of the lowermost Daye limestone, indicated by the first appearance of H.
parvus (Wang and Xia, 2004).
The Chaohu Section situates at the Chaohu City, Anhui
Province. It’s located at the northeast margin of the Yangtze
Platform, much closer to the basin than the Meishan Section
Rapid Carbonate Depositional Changes Following the Permian-Triassic Mass Extinction
169
Figure 3. Microscopic photos of thin-sections from the Zhongzhai Section. (a), (b) and (c) are taken from the lowest 4 cm
while (d) is taken from the uppermost 2 cm. fo. foraminifer; fu. fusulinid; b. brachiopod; m. molluscs; Py. pyrite; q. quartz.
Figure 4. The polished P-T boundary limestones of the studied sections. (a) From the Zhongzhai Section; (b) from the Meishan Section; (c) from the Chaohu Section. The red scale bar is 10 cm in length.
170
Li Tian, Jinnan Tong, David Bottjer, Daoliang Chu, Lei Liang, Huyue Song and Haijun Song
Figure 5. Microscopic photos of thin-sections from the Dajiang Section. (a) and (b) showing the skeleton-enriched packstone,
taken from the uppermost Permian bioclastic limestone; (c) and (d) showing the skeleton-poor wackestone, taken from the
basal Triassic microbialites. al. algae; os. ostracods; cy. calcified cyanobacteria.
Figure 6. Microscopic photos of thin-sections from the Yangou Section. (a) The end-Permian skeleton-enriched grainstone; (b)
the foraminifer-bearing oolitic limestone; (c) and (d) the basal Triassic mollusc-enriched wackestone-packstone. m. Molluscs.
Rapid Carbonate Depositional Changes Following the Permian-Triassic Mass Extinction
171
Figure 7. Microscopic photos of thin-sections from the Meishan Section. (a) and (b) The foraminifer-enriched uppermost
Permian wackestone (Bed 24e); (c), (d) and (e) the mudstone with few skeletons of Bed 27; (f) shows the pyrite-enriched
post-extinction mudstone (Bed 29).
during the P-T transition. The uppermost Permian is composed
of the Dalong Formation, dominated by thin-bedded siliceous
mudstone/limestone while the basal Triassic is the Yinkeng
Formation, comprised of mudstone with muddy limestone (Fig.
2). The lowermost Triassic limestone is well-bioturbated
(ichnofabric index 5; Fig. 4c).
2
RESULTS
Although these six studied sections were located relatively
far apart, the facies and relative water depth are straightforward.
The sandy mudstone dominated Zhongzhai Section was deposited as part of a neritie clastic facies. The occurrence of microbialites at the Dajiang Section and oolites at the Yangou Section
suggest they were on the shallow platform and shoal, respectively. The Meishan Section was on the slope of the north mar-
ginal basin of the Yangtze Platform during deposition of thin
bedded limestones with claystones (Zheng, 2006; Zhang et al.,
1996), while the Xiakou Section and the Chaohu Section were
located off the north margin of the Yangtze Platform in deep
water for siliceous mudstone and claystone deposition. Thus,
it’s clear that the studied sections are distributed from onshore
to offshore, representing different facies deposition (Fig. 2;
Table 1).
The general stratigraphic correlation suggests limestone
was deposited at all six studied sections from onshore to offshore in the P-T transition (Fig. 2). At the Zhongzhai Section,
two distinct limestone beds of medium thickness occur in the P-T
transitional beds while limestone was not deposited in the lower
or upper parts. At the Dajiang Section, the basal Triassic massive
microbialites replaced the uppermost Permian bioclastic
Li Tian, Jinnan Tong, David Bottjer, Daoliang Chu, Lei Liang, Huyue Song and Haijun Song
172
Table 1
The facies analysis of the studying sections
Sedimentary
facies
References
Neritic clastic
Zhang et al., 2014
Thin-medium
bedded limestone, and
thick bedded
or massive
dolomites
Thin-medium
bedded limestone and clay
Extremely shallow platform
Jiang et al., 2014;
Lehrmann et al.,
1998
Shallow carbonate shoal
Tian et al., 2014b;
Sun D Y et al., 2012
Thin-medium
bedded
muddy limestone and clay
Ramp of inner
shelf basin
Zheng, 2006;
Zhang et al., 1996
Tian S F et al.,
2014
Temporal and spatial sediments variations
C. meishanensis H.changxingensis
I. staeschei
Zone
-H. parvus Zone
–I. isarcica
Zone
Sandy mudSandy mudMedium bedded
Sandy mudstone
stone?
limestone and
stone and clay
clay
C. yini Zone
Zhongzhai
Section
Dajiang
Section
Massive
reef like
bioclastic
limestone
Hiatus
Maasive microbiliates
Yangou
Section
Thick bedded bioclastic limestone
Hiatus
Meishan
Section
Mediumthick bedded bioclastic limestone
Thin bedded
clay bed
Thin bedded
oolitic limestone beds,
thin-medium
bedded muddy
limestone
with few clay
beds
Thin bedded
mudstone/shale
and
medium bedded
limestone
Xiakou
Section
Siliceous
mudstone
Mudstone?
Thin bedded
mudstone/shale
and
medium bedded
limestone
Thin bedded
muddy limestone and clay
Basin
Yin et al., 2014;
Wang and Xia,
2004
Chaohu
Section
Siliceous
mudstone
Clay?
Medium-thick
bedded limestone
Thin bedded
muddy limestone and clay
Basin
Yin et al., 2014
Footnote: “?”=not sure.
limestone. At the Yangou Section, there is no remarkable facies
change except for the occurrence of oolites below the PTB. At
the Meishan Section, the boundary limestone (Bed 27) stands out
from the other beds because it is sandwiched by the boundary
clays (Fig. 2e). At the Xiakou Section and the Chaohu Section,
the situations are similar: limestones overlie the siliceous mudstones distinctly and the lowermost limestone bed is thicker than
the overlying beds (Figs. 2f, 2g).
Thus, it’s clear that the boundary limestones cap the uppermost Permian deposits from onshore to offshore. Here, these
limestone beds are designated the Permian-Triassic boundary cap
limestones (PTBCLs). Although the temporal distributions of
PTBCL are restricted from the H. changxingensis Zone to the H.
parvus Zone, the precise correlation (based on conodont zones)
suggests that the bases of PTBCLs are diachronous (Fig. 8).
Besides the occurrences of PTBCLs, no carbonate deposits
have been found within the studied sections during the C. meishanensis Zone. Yin et al. (2014) reviewed twenty three PTB
sections with high resolution conodont zonations in South China,
and there are eighteen sections bearing carbonates in the H. parvus Zone while only three in the C. meishanensis Zone. Thus, the
spatial and temporal variation of carbonate deposits during the
P-T transition can be summed up as 4 phases (Fig. 9): Phase 1)
Pre-extinction, C. yini Zone and earlier, carbonates only occurred
on the platform and slope; Phase 2) C. meishanensis Zone, carbonates rarely occurred; Phase 3) H. changixngensis Zone to H.
parvus Zone, medium-massive bedded carbonates were deposited from onshore to offshore settings; Phase 4) I. staeschei Zone
and later, carbonates disappeared from the neritic clastic facies
and were enriched in mud and thin bedded in other facies.
3
DISCUSSION
The numerous gases vented by the Siberian Trap eruption at
252.2 Ma (Shen S Z et al., 2011; Svensen et al., 2009), were
hypothesized to have led to global warming (Joachmiski et al.,
2012; Sun Y D et al., 2012), strong terrestrial weathering (Algeo
et al., 2011; Algeo and Twichett, 2010) and euxinia/anoxia
(Kaiho et al., 2006; Grice et al., 2005; Wignall and Twichett,
Rapid Carbonate Depositional Changes Following the Permian-Triassic Mass Extinction
173
Figure 8. Lithological and conodont zonation correlation for the studied sections. This figure is modified from Yin et al. (2014)
with additional data. The conodont zones are correlated with published data (Zhongzhai Section: Zhang et al., 2014; Dajiang
Section: Jiang et al., 2014; Yangou Section: Sun D Y et al., 2012; Meishan Section: Jiang et al., 2007; Xiakou Section: Wang
and Xia, 2004).
1996), contributing to the P-T mass extinction and the delayed
Early Triassic recovery in its aftermath (Pietsch and Bottjer, 2014;
Song H J et al., 2014, 2012a, b; Tian et al., 2014a; Sun Y D et al.,
2012; Algeo et al., 2011; Bottjer et al., 2008).
Along with these hypotheses, ocean acidification also is a
reasonable consequence of rapid volcanic eruption (Kump et al.,
2009; Wignall, 2001). The preferential extinction of heavier
calcified fossils was demonstrated as a biotic response to ocean
acidification (Knoll et al., 2007, 1996) while extensive erosion
truncation surfaces within the subaerial carbonate P-T boundary successions were documented as the sedimentary response
(Payne et al., 2007). However, truncation surfaces and strata
loss between Permian bioclastic limestone and basal Triassic
microbialite also could be attributed to the exposure of shallow
carbonates during the end-Permian regression (Yin et al., 2014;
Kershaw et al., 2012; Collin et al., 2009; Wignall et al., 2009).
The controversial explanations for these sedimentary changes
have inspired us to synthesize these scenarios and propose
other hypotheses for fluctuations of the carbonate deposition
during this time.
3.1 Oversaturation of Carbonates Following Ocean Acidification?
During volcanic eruption extinction events, ocean acidification is an essential environmental effect of the increase of
CO2 or other acid hydrosoluble gases (Greene et al., 2012; Algeo et al., 2011; Wignall, 2001). The huge amount of CO2 released by the gigantic Siberian Trap eruption would have unbalanced the pre-extinction climate and ocean chemistry, contributing to the P-T mass extinction (Hinojosa et al., 2012;
Payne and Clapham, 2012; Payne et al., 2010, 2007; Knoll et
al., 2007). In this scenario, following the intense injection of
oceanic CO2, CaCO3 dissolution would overcome precipitation
because of the onset of ΤCO2 (Equation 1) increase and constant
cALK (Equation 2) input (Kump et al., 2009). However this
acidified ocean would be reversed by reduced carbon burial
(Greene et al., 2012; Rampino and Caldeira, 2005) and alkaline
input, since sustainable increase of atmospheric CO2 would
subsequently enhance terrestrial weathering (Greene et al.,
2012; Algeo et al., 2011; Kump et al., 2009). Cui et al. (2013)
modeled a similar carbonate saturation dynamic, showing that
the ocean might be oversaturated in 20 Ka. after the onset of
concurrent mass extinction and volcanic eruption. This duration
174
Li Tian, Jinnan Tong, David Bottjer, Daoliang Chu, Lei Liang, Huyue Song and Haijun Song
Figure 9. Temporal profiles of palaeoceanographic conditions during the P-T transition. Species richness is from Song H J et
al. (2013); dating data are from Burgess et al. (2014). The carbonate deposition ranges are summarized by this study while
the ocean redox conditions are after Song H J et al. (2012a) and Li et al. (2010).
could be less than 10 Ka., as calculated by Kump et al. (2009).
The abrupt increase of CO2 does cause a drop in surface
seawater PH, but the significant dissolution of carbonate only
happens during the rapid injection, of which the duration is less
than 10 000 years (Hönisch et al., 2012). Although the total
amount of the CO2 increase is huge (Svensen et al., 2009), the
average annual carbon release is even much smaller than fossil
fuel consumption in modern times (Hönisch et al., 2012; Kump
et al., 2009). Ca isotopes in bulk samples and conodonts
(Hinojosa et al., 2012; Payne et al., 2010) as well as 187Re/188Os
(Georgiev et al., 2011) were proposed as geochemical evidences for end-Permian ocean acidification, but none of these
is a direct proxy for sea water PH. Weathering also could contribute to these geochemical anomalies (Kershaw et al., 2012).
Considering the P-T extinction interval is just 0.061 Ma. (Burgess et al., 2014), the over Ma. residence time of Ca isotopes
invalidates its efficiency as an indicator of ocean acidification
during the rapid P-T transition.
Payne et al. (2010) documented Ca isotope evidence for
the ocean acidification hypothesis, which can account for the
greater effect on heavily calcified animals (Knoll et al., 2007,
1996) and extensive submarine erosional surfaces (Payne et al.,
2007) in the extinction interval, as well as the additional oolites
and microbialites in the post-extinction carbonate supersaturated oceans (Payne et al. 2007; Baud et al., 2005). However,
these observed sedimentary phenomenon also can be attributed
to Mg/Ca seawater chemistry anomaly and rapid regression
(Hönisch et al., 2012; Kershaw et al., 2012; Collin et al., 2009;
Kump et al., 2009; Wignall et al., 2009). The saturation rate in
the post-extinction (over 20 Ka.) is just 0.2 larger than the
pre-extinction and no overshoot increase of pH is seen in the
modeling (Cui et al., 2013). Although the weathering of CaSiO3
and CaCO3 presented by Kump et al. (2009) also just increases
cALK by 2 and 2, as well as ΤCO2 by 2 and 1, respectively (Table
1), such equivalent increases in both cALK of ΤCO2 won’t cause
significant change of saturation condition. Meanwhile, Kump et
al. (2009) misused the late Spathian oolites and microbialites (5
Ma. later than the P-T mass extinction) as indicators of the
oversaturated ocean following the P-T mass extinction. Recent
review of oolites suggests that oolites began global blooming
from the H. changxingensis Zone (Li et al., 2014), which is a
typical post major extinction conodont zone (Metcalfe et al.,
2007). Thus, seawater could have been carbonate oversaturated
after the major phase of P-T mass extinction, as indicated by
oolites and microbialtes, but the significant ocean acidification
during or at the onset of extinction still needs further direct
evidence to verify.
ΤCO2=[CO2]+[HCO3-]+[CO32-]
cALK=[HCO3-]+2[CO32-]+[B(OH)4-]+[OH-]–[H+]
3.2
(1)
(2)
Subsequent Sedimentary Response to Upwelling?
The end-Permian to Early Triassic oceans were thought to
be comparative to the Precambrian and the Black Sea in modern times (Woods et al., 1999; Knoll et al., 1996; Grotzinger
and Knoll, 1995). The OMZ expanded in such stratified oceans
(Algeo et al., 2011; Knoll et al., 1996), a scenario supported by
numerous direct mineral and geochemistry evidence, such as
framboidal pyrites and trace elements (Tian et al., 2014a; He et
al., 2013; Song H J et al., 2012b; Bond and Wignall, 2010; Liao
Rapid Carbonate Depositional Changes Following the Permian-Triassic Mass Extinction
et al., 2010; Shen et al., 2007). The strong sulfate reduction is
one of the most distinct characteristics that differ from normal
open oceans, as demonstrated by the perturbations of sulfur
isotopes (Song H Y et al., 2014; Luo et al., 2010). Furthermore,
photosynthesis in surface water and sulfur reduction in deep
water leads to carbon isotopes that form a large gradient between shallow and deep (Song H Y et al., 2013; Song H J et al.,
2012a; Meyer et al., 2011; Knoll et al., 1996). Since sulfate
reduction could increase the alkalinity (Table 1), the mixing of
deep water with shallow water via upwelling was demonstrated
to have contributed to carbonate precipitation on the shelf of
the Black Sea (Kempe, 1990). The physical and biological CO2
degassing from the upwelled alkaline water could lead to carbonate oversaturation. This process was thought to have happened in the Permian and Early Triassic oceans, causing seafloor fan precipitates and growth of microbialites (Kershaw et
al., 2007; Woods et al., 1999; Grotzinger and Knoll, 1995). The
sulfur isotope signal recorded upwelling in the end-Permian
pre-extinction interval rather than the post-extinction (Shen Y A
et al., 2011). The carbon isotope gradient also shows a negative
shallow-deep value (indicating mixing or upwelling) at the
end-Permian pre-extinction time while it’s reversed to be positive (indicating stagnation) after the extinction (Song H J et al.,
2012a). Based on these observations, the upwelling was prior to
the extinction event while the post-extinction oolites and microbialites might be its subsequent sedimentary response.
However, the loopholes of this scenario are from the pCO2
and sulfur oxygenation. The high pCO2 would elevate oxygen
and oxygenate the sulfur during mixing, which is a reverse
process of sulfate reduction (Kempe, 1990), in contrast to the
observed sulfur isotope perturbations (Song H Y et al., 2014;
Luo et al., 2010). The sulfur reduction decreases in the cALK,
leading to the dissolution of carbonate rather than precipitation
(Table 1). Meanwhile, the upwelling hypothesis is also unconstrained with climate driving ocean circulation models. Algeo
et al. (2011) demonstrated that the warming events correspond
with stagnation while the cooling events correspond with upwelling, but there is no temperature decrease of surface seawater that has been detected in the P-T transition (Joachmiski et
al., 2012). In addition, the upwelling also will elevate productivity by providing nutrients to shallow marine environments
(Coale et al., 1996; Pace et al., 1987; Suess, 1980), leading to a
significant increase in carbon isotopes, which hasn’t been
documented in the P-T extinction interval neither (Song H J et
al., 2012a; Kershaw et al., 2007; Xie et al., 2007). The atmosphere-ocean modeling results also suggest that there was no
significant upwelling in the South China Block (Kidder and
Worsley, 2004). Therefore, no matter the theoretical speculations and the observed evidences, neither of them can support
the upwelling hypothesis.
3.3 Anaerobic Seawater Oxygenation: A Potential Geochemical and Biogeological Process During the P-T Transition
In modern oceans, phytoplankton growth is limited by the
availability of iron, so that efficient iron supply would rise the
productivity in surface water significantly (Takeda, 1998; Coale
et al., 1996; De Baar et al., 1995). Similarly, the enhanced iron
175
supply also might have played a key role in microbial blooms
during the P-T interval. The Sr isotopes and other geochemical
indexes all indicate an enhanced weathering and terrestrial
input during the P-T mass extinction and its aftermath (Sedlacek et al., 2014; Algeo et al., 2011; Algeo and Twichett, 2010;
Korte et al., 2003), associated with elevated iron input. Meanwhile, the large perturbations of sulfur isotopes (Song et al.,
2014; Luo et al., 2010) suggests strong sulfur reduction, which
could produce abundant H2S in shallow water, so that the mixing of iron and sulfur in shallow water would react as pyrite
precipitation (Table 2). In this reaction, the ΔcALK is 4 while the
ΔΤCO2 is 0. The imbalanced increase ratio of ΔcALK to ΔΤCO2
provides the kinetic for the supersaturation of carbonates. The
observed calcified microbes within microbilates (Yang et al.,
2011; Wang et al., 2005) are modest biotic evidences for increased alkalinity. The appearance of abundant pyrite in the
shallow water microbialites (He et al., 2013; Bond and Wignall,
2010; Liao et al., 2010) confirms another product in pyrite precipitation, FeS2. Although the sulfur is oxygenated, it’s not a
complete sulfur oxygenation otherwise it would balance the
low sulfate after the strong sulfate reduction. In addition, if the
large amount of CH4, which was injected by the Siberian Trap
eruption, was oxygenated into CH2O in seawater by microbial
activities, the geochemical process will be like anaerobic organism oxygenation in Table 2, in which the ΔcALK is 8 while
ΔΤCO2 is 1, leading to dramatic carbonate supersaturation.
Looking into these geochemical processes, the irons are
not oxygenated at all but sulfur and methanol are oxygenated.
The iron hasn’t been oxygenated into hematite. There is no
available literature about the significant increase of hematite in
marine sediments during the P-T transition. He et al. (2013)
proposed there was an oxygenation event above the P-T
boundary because of the increased diameter of framboidal pyrites. The occurrence of bioturbation in the PTBCLs of Meishan and Chaohu suggest a dysoxia-oxic environment (Figs.
4b, 4c). The distinct decrease of the carbon isotope vertical
gradient in the transitional beds also implies a rapid ventilation
of seawater (Song H J et al., 2012a). However, no O2 is involved in these reactions except for sulfur reduction (Table 2).
Given the perspective of oxygenation, these geochemical processes suggest an anaerobic oxygenation rather than a complete
oxygenation, implying a low O2 concentration during the P-T
transition. In modern times or normal environments, the seawater above fair weather wave base contains sufficient O2,
which is being aerated from the atmosphere. Berner (2006)
modeled out a drop of atmospheric O2 concentration from 30 to
15% in the latest Permian and its aftermath. A number of studies show marine euxinia-anoxia is widely spread in oceans
(Wignall and Twichett, 2002, 1996) of this time, especially
during the onset of the mass extinction (Shen et al., 2007; Grice
et al., 2005), but the redox conditions were unstable and in
large perturbation during Early Triassic (Tian et al., 2014a;
Song H J et al., 2012a).
3.4
Implications for P-T Ecosystems
Although we have observed the PTBCLs from onshore to
offshore settings at the onset of the main extinction, the low
sedimentation rates and pyrite poor deposition in other facies
Li Tian, Jinnan Tong, David Bottjer, Daoliang Chu, Lei Liang, Huyue Song and Haijun Song
176
Table 2
Geobiological process
Simplified key geobiological processes and effects on carbonate saturation
Carbonate chemistry
Carbonates weathering
Silica weathering
Sulfate reduction
Sulfur oxygenation
CO2+H2O+CaCO3=2HCO3-+Ca2+
2CO2+H2O+CaSiO3=2HCO3-+Ca2++SiO2
2CH2O+SO42-=2HCO3-+H++HSH++HS-+2O2+2HCO3-=SO42-+2H2O+2CO2
Carbonate precipitation
Pyrite precipitation
Anaerobic organism
oxygenation
Ca2++2HCO3-=CO2+CaCO3+H2O
2HS-+2FeOOH+4H+=FeS2+Fe2+4H2O
4FeOOH+CH2O+8H+=4Fe2++CO2+7H2O
ΔcALK
(mol)
2
2
1
-1
ΔΤCO2
(mol)
1
0
0
0
Carbonate
saturation
+
+
+
-
-2
4
8
-1
0
1
++
+++
Footnotes: -. decrease; +. slight increase; ++. considerably increased; +++. dramatically increased.
suggest that geobiolgical processes varied among facies. The
anaerobic oxygenation process we proposed above could just
have played a key role in the carbonate supersaturation in the
shallow carbonate platform and shoal facies, contributing to the
rapid sedimentation rate. Algeo and Twichett (2010) documented remarkably rapid sedimentation in the Early Triassic at
the Meishan and the Chaohu sections, by a 7X increase from
the Upper Permian. However, the sedimentation rate of the
P-Tr extinction transition beds (beds 25–28) is much lower than
the Upper Permian and the Early Triassic at the Meishan Section (Burgess et al., 2014).
The Xiakou and Chaohu sections were deeper than the
Meishan Section at that time, but the limestone deposits of
phase 2 are thicker and better oxygenated, indicated by more
extensive bioturbation (Fig. 4c). It seems that the ocean was
reversed: the basinal water was better oxygenated than slope
water. This corresponds more closely to the OMZ expanding
model (Algeo et al., 2011) rather than the ocean stagnation
model (Isozaki, 1997), because in a modern stagnated ocean,
the Black Sea, the dissolved O2 decreases with water depth
(Luther et al., 1991). It’s impossible that deeper water is better
oxygenated than slope water in a stagnant ocean. Therefore,
regarding the expanded OMZ as the oceanic model of that time,
the co-occurring H2S probably contributed to slower sedimentation within the OMZ, since the H2S rich water is more acidic
than normal water. This model is similar to the “sandwiched
sea” in the Ediacaran (Li et al., 2010): there is an H2S edge at
the stratified ocean shelf, except the bottom water is oxygenated during the P-T transition rather than ferruginous in the
Precambrian. Meanwhile, in the H. parvus Zone, oxygenation
reduced the sulfur and prevented the formation of small framboidal pyrites (Bond and Wignall, 2010; Shen et al., 2007).
This process would cause the frequent under-saturation of carbonates, leading to the hardground structure within the PTBCL
(Fig. 7e) as well as reduced sedimentation rate of the Meishan
Section and lack of carbonate at the Shangsi Section.
Looking into the time line, it’s obvious that PTBCL began
deposition from the H. changxingensis Zone at the Yangou and
the Meishan sections, but PTBCLs didn’t occur until the H.
parvus Zone at other sections. Considering that eustatic
changes might expose the shallow platform above sea level
during the C. meishanensis to the H. changxingensis Zones
(Yin et al., 2014), it’s possible that the carbonate supersatura-
tion happened earlier in shallow water settings (but to the
neritic clastic facies) than in deep water. Joachmiski et al.
(2012) recorded a double warming during the P-T transition at
the Meishan Section, but there is no significant change within
Bed 27. The relatively stable temperature could be accounted
for by the pause of catastrophic events, although the rock record shows a diachronous sedimentary response for the spatial
geochemical and geobiological carbonate variations. The
shrinking of the OMZ, related to the H2S edge, and the enhanced weathering input in phase 2, would not just have contributed to the formation of PTBCLs. This rapid break from the
disastrous aftermath of the first episode of the P-T mass extinction could also have provided the survival species a relatively
adaptable zone to inhabit, demonstrated by the abundant ostracods and gastropods at the Dajiang and Yangou sections, the
foraminifers at the Zhongzhai and Meishan sections (Song H J
et al., 2013) and intensive bioturbation at the Chaohu Section,
within PTBCLs.
4
CONCLUSION
The global spread basal Triassic microbialites is thought to
be a sedimentary response to the P-T mass extinction, along
with other unusual carbonate sediments, such as stromatolites,
oolites and sea floor fan precipitates (Knoll et al., 2007; Pruss
et al., 2006). If the Early Triassic oceans really stratified according to the model of Knoll et al. (1996), photosynthesis and
sulfate reduction might change the carbonate saturation of the
shallow photic zone and the deep anoxic water, respectively.
However, the extended carbonate depositional zone in the earliest Triassic we observed suggests that photosynthesis was
limited. Even if there was an oxygenation event, it was minor
and lacked sufficient O2. The strong bioturbation in the deep
water section also reveals that Early Triassic oceans were not as
stratified as previously understood (Isozaki, 1997; Knoll et al.,
1996). The spatial variations of the PTBCLs reveal that carbonate deposition varied within facies and might have changed
with different biological and geochemical processes.
All these evidences and theoretical analysis imply that
rapid carbonate deposition changes were affected by enhanced
terrestrial weathering input, abnormal ocean circulation and
various geobiological processes, associated with metazoan
extinction and the occurrences of unusual sediments. The ocean
geochemistry could be reversed as “Strangelove Ocean”, in
Rapid Carbonate Depositional Changes Following the Permian-Triassic Mass Extinction
which the primary productivity was replaced by the terrestrial
input as the major contributor of the carbonate factory, during
the Permian-Triassic transition (Rampino and Caldeira, 2005;
Kump, 1991). Zeebe and Westbroek (2003) demonstrated the
carbonate is critical supersaturated in the “Strangelove Ocean”
dominated by inorganic precipitation. Meanwhile, the rapid
carbonate deposition changes presented in this study reveal that
various geochemical and geobiological processes took over
regular biomineralization and carbonate deposition due to the
low O2 and high CO2, responding to the large volcanic eruptions.
ACKNOWLEDGMENTS
This study was supported by the “973 Program” (No.
2011CB808800), the National Natural Science Foundation of
China (Nos. 41172312, 41272372, 41302010, 41402302), the
State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences (No. GKZ14Y663) and
the “111 Project” (No. B08030). We thank Will Berelson and
Frank Corsetti in USC for the teaching fundamental carbonate
geochemistry, as well as constructive discussion with Fei Li.
We also thank the editor and two anonymous reviewers for the
helpful comments.
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