ORIGINAL RESEARCH
published: 24 June 2021
doi: 10.3389/feart.2021.651224
Hg Isotopes and Enhanced Hg
Concentration in the Meishan and
Guryul Ravine Successions: Proxies
for Volcanism Across the
Permian-Triassic Boundary
Edited by:
Jun Shen,
China University of Geosciences
Wuhan, China
Reviewed by:
Hengye Wei,
Southwest Petroleum University,
China
Shane Schoepfer,
Western Carolina University,
United States
Runsheng Yin,
Institute of Geochemistry, Chinese
Academy of Sciences (CAS), China
*Correspondence:
Alcides Nóbrega Sial
alcides.sial@ufpe.br
Jiubin Chen
chenjiub@hotmail.com;
Specialty section:
This article was submitted to
Geochemistry,
a section of the journal
Frontiers in Earth Science
Received: 08 January 2021
Accepted: 31 May 2021
Published: 24 June 2021
Citation:
Sial AN, Chen J, Korte C, Pandit MK,
Spangenberg JE, Silva-Tamayo JC,
Lacerda LDde, Ferreira VP,
Barbosa JA, Gaucher C, Pereira NS
and Riedel PR (2021) Hg Isotopes and
Enhanced Hg Concentration in the
Meishan and Guryul Ravine
Successions: Proxies for Volcanism
Across the PermianTriassic Boundary.
Front. Earth Sci. 9:651224.
doi: 10.3389/feart.2021.651224
Alcides Nóbrega Sial 1*, Jiubin Chen 2*, Christoph Korte 3, Manoj Kumar Pandit 4,
Jorge E. Spangenberg 5, Juan Carlos Silva-Tamayo 6, Luiz Drude de Lacerda 7,
Valderez Pinto Ferreira 1, José Antônio Barbosa 1, Claudio Gaucher 8, Natan Silva Pereira 9
and Paulo Ricardo Riedel 1
1
Department of Geology, NEG–LABISE, UFPE, Recife, Brazil, 2School of Earth System Science, Tianjin University, Nankai, China,
Nordic Center for Earth Evolution (NordCEE), Department Geosciences and Natural Resource Management, University of
Copenhagen, Copenhagen, Denmark, 4Department of Geology, University of Rajasthan, Jaipur, India, 5Institute of Earth Surface
Dynamics, University of Lausanne, Lausanne, Switzerland, 6One-Health Research Group, Testlab Laboratorio Analisis Alimentos
y Aguas SAS, Medellin, Colombia, 7LABOMAR, Institute of Marine Sciences, UFC, Fortaleza, Brazil, 8Instituto de Ciencias
Geológicas, Facultad de Ciencias, Universidad de La República, Montevideo, Uruguay, 9Department of Biology, State University
of Bahia, Campus VIII, Paulo Afonso, Brazil
3
High-resolution organic carbon isotope (δ 13C), Hg concentration and Hg isotopes curves
are presented for the Permian-Triassic boundary (PTB) sections at Guryul Ravine (India)
and Meishan D (China). The total organic carbon (TOC)-normalized Hg concentrations
reveal more intense environmental changes at the Latest Permian Mass Extinction (LPME)
and the earliest Triassic Mass Extinction (ETME) horizons coinciding with major δ 13C shifts.
To highlight palaeoredox conditions we used redox-sensitive elements and Rare Earth
Element distribution. At Meishan, three Hg/TOC spikes (I, II, and III) are observed. Spike I
remains after normalization by total aluminum (Al), but disappears when normalized by total
sulfur (TS). Spike III, at the base of Bed 26, corresponds with excursions in the Hg/TS and
Hg/Al curves, indicating a change in paleoredox conditions from anoxic/euxinic in the
framboidal pyrite-bearing sediments (Bed 26) to oxygenated sediments (Bed 27). At
Guryul Ravine, four Hg/TOC spikes were observed: a clear spike I in Bed 46, spike II at the
base of the framboidal pyrite-rich Bed 49, spike III at the PTB, and spike IV at the LPME
horizon. Some of these Hg/TOC spikes disappear when TS or Al normalization is applied.
The spike I remains in the Hg/TS and Hg/Al curves (oxic conditions), spike II only in the Hg/
TS curve (anoxic/euxinic), and spikes III and IV only in Hg/Al curves (oxic). In both sections,
Hg deposition was organic-matter bound, the role of sulfides being minor and locally
restricted to framboidal pyrite-bearing horizons. Positive mass-independent fractionation
(MIF) for Hg odd isotopes (odd-MIF) was observed in pre-LPME samples, negative values
in the LPME–PTB interval, and positive values above the ETME horizon. Most Hg-isotope
patterns are probably controlled by the bathymetry of atmospheric Hg-bearing deposits.
The source of Hg can be attributed to the Siberian Traps Large Igneous Province (STLIP).
In the LPME-PTB interval, a complex of STLIP sills (Stage 2) intruded coal-bearing
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Sial et al.
Hg Chemostratigraphy Permian-Triassic Boundary
sediments. The negative δ 202Hg, the mercury odd-MIF Δ201Hg patterns, and the
Δ199Hg–Hg plot in both sections are compatible with volcanic mercury deposition. Our
study shows the strength of Hg/TOC ratios as paleoenvironmental proxy and as a tool for
stratigraphic correlation.
Keywords: permian, triassic boundary, meishan GSSP, Hg isotopes, chemostratigraphy, mass extinction
INTRODUCTION
Thibodeau and Bergquist, 2017; Grasby et al., 2017; Grasby
et al., 2019; Gong et al., 2017; Wang et al., 2018; Them et al.,
2019; Shen et al., 2019a,b,c; Shen et al., 2020).
Major Phanerozoic mass extinctions seem to be coeval with the
volcanic activity of Large Igneous Provinces (LIPs; e.g., Courtillot,
1994), a correlation that was observed in several studies (e.g.,
Courtillot et al., 1999; Wignall, 2001; Courtillot and Renne, 2003;
Kravchinsky, 2012; Bond and Wignall, 2014; Bond and Grasby,
2017). The major challenges in demonstrating the linkage
between the five major Phanerozoic mass extinctions (Raup
and Sepkoski, 1982) and LIP eruptions reside in the lack of
accurate radiometric dating of LIP volcanic rocks and of
sedimentary rocks, which recorded abrupt biological and
environmental changes leading to mass extinctions (Grasby
et al., 2019). However, recent advances on the radiometric
dating of LIP volcanism (e.g., Burgess et al., 2014; Burgess and
Bowring, 2015; Renne et al., 2015; Schoene et al., 2015; Schoene
et al., 2019; Sprain et al., 2019) have strengthened the connection
between LIP volcanism and mass extinction.
Recent studies show that Mercury (Hg) in sedimentary
successions is a robust proxy of massive volcanism in periods
of extreme environmental turnover (e.g., Sial et al., 2010, Sial
et al., 2013, Sial et al., 2014, Sial et al., 2016, Sial et al., 2019, Sial
et al., 2020a, Sial et al., 2020b; Nascimento-Silva et al., 2011;
Nascimento-Silva et al., 2013; Sanei et al., 2012; Grasby et al.,
2015a; Grasby et al., 2017; Grasby et al., 2019; Grasby et al., 2020;
Percival et al., 2015; Percival et al., 2017; Percival et al., 2018;
Adatte et al., 2015; Font et al., 2016; Jones et al., 2017; Gong et al.,
2017; Thibodeau et al., 2016; Thibodeau and Bergquist, 2017;
Charbonnier et al., 2017; Sabatino et al., 2018; Burger et al., 2019;
Meyer et al., 2019; Shen et al., 2019a,b,c, Shen et al., 2020; Them
et al., 2019; Keller et al., 2020; Georgiev et al., 2020). Mercury
enrichments across some chronological boundaries, expressed as
an increase in the Hg to total organic carbon ratio (Hg/TOC),
may represent true volcanogenic Hg loading to the environment
(e.g., Sanei et al., 2012; Grasby et al., 2013; Percival et al., 2015;
references therein). Therefore, the search for Hg/TOC spikes
across major chronostratigraphic boundaries has been expanded
to the whole Phanerozoic Eon (e.g., Grasby et al., 2013; Grasby
et al., 2015b; Grasby et al., 2017; Grasby et al., 2019; Grasby et al.,
2020; Bond and Grasby, 2017; Gong et al., 2017; Jones et al., 2017;
Percival et al., 2015, 2017, 2018; Burger et al., 2019; Korte et al.,
2019; Shen et al., 2019a,b,c, 2020; Sial et al., 2016; Sial et al., 2019;
Sial et al., 2020a; Sial et al., 2020b; Faggetter et al., 2019; Kwon
et al., 2019). Despite the growing application of Hg
chemostratigraphy as a proxy for LIP activity in sedimentary
successions, only a few studies have used Hg isotopes as an
additional tool to demonstrate the volcanogenic origin of Hg
enrichments (e.g., Sial et al., 2014; Sial et al., 2016; Sial et al., 2019;
Sial et al., 2020a; Sial et al., 2020b; Thibodeau et al., 2016;
Frontiers in Earth Science | www.frontiersin.org
Causes of the Permian–Triassic Mass
Extinction
The most severe biodiversity decline of the eukaryotic biota in
Earth history took place at 251.9 Ma (e.g., Erwin et al., 2002;
Benton, 2003; Chen and Benton, 2012; Burgess et al., 2014;
Burgess and Bowring, 2015; Renne et al., 2015). This mass
extinction caused the death of over 90% of all marine and
about 70% of all terrestrial species (e.g., Erwin, 2006). The
event is known as the P–Tr extinction, the P–T extinction,
End-Permian Mass Extinction (EPME), Latest Permian Mass
Extinction (LPME), and informally the “Great Dying”. Several
causes have been invoked to explain this mass extinction and,
among them, shallow oceanic water anoxia, bolide impact, and
volcanism seem to be the main ones.
A widespread episode of oceanic photic-zone anoxia has been
widely documented in PTB sections through the presence of
framboidal pyrite, regarded as evidence of anoxia/euxinia (Bond
and Wignall, 2010). Notable among these are the sections in
Greenland and Sosio Valley (Wignall and Twitchett, 2002),
southwestern Japan (Isozaki, 1997), Italy and Slovenia
(Wignall and Twitchett, 1996), Kashmir in northern India
(Wignall et al., 2005; Huang Y. et al., 2019) and British
Columbia, Canada (Wignall and Newton, 2003; Bond and
Wignall, 2010). A worldwide dysoxic marine event during the
P–Tr transition, suggested by Shen et al. (2007) in a study of the
size/distribution of framboidal pyrite at the Meishan section,
seems to reinforce a close relationship between oxygen
availability and the marine LPME. This episode of oceanic
photic anoxia may have lasted 10 Myr, and is regarded as one
prime cause of the P–Tr marine biotic crisis (Wignall and Hallam,
1992; Isozaki, 1997; Krystyn et al., 2003; Lehrmann et al., 2003;
Wignall et al., 2005; Shen et al., 2007; Song et al., 2014; Kaiho
et al., 2016; Kumar et al., 2017; Huang Y. et al., 2019) and named
“superanoxic event” (Isozaki, 1997; Grice et al., 2005).
Wignall and Newton (2003) proposed that the LPME is
diachronous by ≥ 0.5 Myr with the late Changxingian
extinction in Panthalassa, coinciding with the diversity
increase that led to the migration of warm-water taxa into the
high southerly paleolatitudinal regions of the Neotethys Ocean.
These authors suggest that during the Changxingian extinction in
Panthalassa, there was a decline in the seafloor oxygen availability
(absence of bioturbation and presence of framboidal pyrite),
followed by euxinic conditions in the latest Changxingian and
Early Triassic (e.g., Ursula Creek, British Columbia). However, in
2
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Sial et al.
Hg Chemostratigraphy Permian-Triassic Boundary
FIGURE 1 | (A) Late Permian paleogeographic map for the Meishan D (GSSP) section, China (based on Wu et al., 1983; Cao et al., 2010; Li and Jones, 2017). (B)
Guryul Ravine section, Kashmir (modified from Tewari et al., 2015) and (C) global palaeogeographical reconstruction for the Permian–Triassic transition and geographical
location of Meishan GSSP and Guryul Ravine sections (modified from Sial et al., 2020a). (A) OPC open platform carbonate, CARL calcareous algal reef limestone,
CRCL carbonate ramp cherty limestone, DWM deep-water mudstone, LA land area. (B) HA Holocene alluvium, TS Triassic strata, PS Permian strata,
PV Panjal Volcanics.
characterized by low pyrite sulfur isotope ratios (δ 34S) and low
carbon isotope ratios of carbonates (δ 13Ccarb), with the first
pyritic horizon coinciding with the LPME, suggest upwelling
of toxic deep waters as a major process at the P–Tr transition
(Algeo et al., 2008).
Anoxia responsible for the global PTB biodiversity crisis may
have been caused by global warming linked to LIP eruptions
(Bond et al., 2020). The toxic effect of volatile-rich mafic lavas or
sill-type intrusions depends on the size of the igneous province,
geological setting of the host region, magma plumbing system,
and eruption dynamics that control the magnitude and
composition of the thermogenic outgassing (Racki, 2020).
Therefore, it was suggested that the extensive volcanism of the
Siberian Traps Large Igneous Province (STLIP) was likely the
main trigger of the PTB biodiversity crisis (e.g., Campbell et al.,
1992; Yin et al., 1992, 2007; Renne et al., 1995; Bowring et al.,
1998; Courtillot et al., 1999; Lo et al., 2002; Kamo et al., 2003;
Grard et al., 2005; Reichow et al., 2009; Korte et al., 2010; Burger
et al., 2019). Mercury anomalies have been described globally
from several Permian-Triassic sections, implying extensive
volcanism during the Permian-Triassic transition (Grasby
et al., 2017; Shen et al., 2019b; Sial et al., 2020a; Sial et al., 2020b).
The atmospheric CO2 partial pressure largely increased during
STLIP eruptions at the end-Permian (252 Ma ago). This increase
may have triggered ocean acidification across the LPME, as
suggested by calcium and boron isotopes (Payne et al., 2010;
Hinojosa et al., 2012; Clarkson et al., 2015; Silva-Tamayo et al.,
2018) and U isotopes (Lau et al., 2016), stressing the ocean biota.
This possibility was evaluated by Kershaw et al. (2012) by
comparing past and modern ocean acidifications, the latter
one raised by the rapid current increase in atmospheric CO2.
However, these authors concluded that large increases in CO2 in
the Selong PTB section in South Tibet, a recorded regression and
erosion in the mid-Changxingian was followed by a deepening in
the late Changxingian and across the P–Tr transition. The latter is
marked by an increase in faunal diversity, suggesting warming
conditions without suffering marine extinction. Permian
remnants perished only in the disoxic, thinly bedded and
pyrite-rich limestone level of the late Griesbachian (Wignall
and Newton, 2003).
During the P–Tr transition, equatorial and higherpaleolatitude settings displayed a zonation with severe anoxia
prevailing in the former setting and less intense and short-lived
anoxia recorded in the latter (Bond and Wignall, 2010). These
authors inferred from the analysis of the size of framboidal pyrite
that: a) intense anoxia with euxinia developed throughout the
P–Tr interval from the Boreal and the Neotethys oceans, and
dysoxia, in relatively shallow-water settings above the storm wave
base; b) dysoxic conditions with rare euxinia, in contrast, widely
developed over a range of water depths (including shallow ones),
at equatorial paleolatitudes; c) a complex and unstable redox
history was recorded in western and eastern Tethyan locations,
with the ventilation of anoxic environments (Hindeodus
praeparvus Zone) during the P–Tr interval (H. changxingensis
to H. parvus zones), returning later to oxygen-poor environments
(Isarcicella isarcica Zone). The extent and timing of anoxia in
Gondwana during the P–Tr transition is poorly documented
(Huang Y. et al., 2019). Kaiho et al. (2016) proposed that at low
latitudes, an increase in temperature triggered soil erosion and a
gradual change from a well-mixed oxic to a stratified euxinic
ocean before the LPME. At the ETME, about 60 kyr later, anoxia
in nearshore surface waters and anoxia/euxinia in shallowintermediate waters lasted for almost 1 Myr. In a marine PTB
section at Nhi Tao, Vietnam, a series of pyritic horizons
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TABLE 1 | Trace element (ppm) and elemental ratios (including rare-earth elements) for the Meishan and Guryul Ravine sections.
Meishan Section
Period
Fm.
Triassic
Yinkeng
Permian
Changxing
Bed
30
29D
29C
27D
27C
27B
27A
26
26
25
24E
4
24D
24C
24B
Sample
Height(cm)
Meishan 41
Meishan 40
Meishan 38
Meishan 36
Meishan 33
Meishan 29
Meishan 26
Meishan 24
Meishan 51B
Meishan 51A
Meishan 50
Blw Mei. 50
Meishan 23
Meishan 20
Meishan 15
Meishan 11
Meishan 8
68.7
56.5
36.5
25.5
20.5
15.0
10.5
6.5
1.9
0.8
−0.1
−0.2
−3.5
−7.5
−22.5
−68.0
−88.0
Detrital Supplay
Productivity Proxies
Pal
Th/
U
Cr/
Th
Zr/
Cr
Zr
REE
+Y
Y/Ho
Zr/Hf
CuEF
NiEF
Ni/
Co
Ba/
Al(x10
Mn/
Fe
Pb
V/
Cr
V/(V
+ Ni)
VEF
MoEF
4.47
3.49
3.84
1.59
1.33
0.75
1.84
2.88
8.86
6.24
1.14
0.94
0.14
0.23
0.09
0.18
0.35
4.56
4.72
5.05
4.04
4.15
4.43
4.25
4.57
0.07
0.76
1.22
10.42
11.74
9.38
12.53
15.38
28.30
2.44
2.24
2.02
2.30
2.19
2.20
2.15
1.99
90.67
9.52
6.06
1.03
0.76
0.83
0.69
0.37
0.59
101.75
78.30
55.93
54.57
49.08
58.86
122.38
142.44
372.52
282.88
60.23
266.28
4.19
5.67
3.74
1.12
27.39
158.16
139.73
121.29
167.46
167.02
227.59
257.15
219.83
409.14
190.44
142.64
453.07
22.02
29.05
10.77
7.14
34.42
31.09
32.12
33.11
33.35
35.09
36.30
31.99
29.16
24.85
24.60
23.07
24.82
46.51
37.88
42.55
42.65
38.47
35.85
36.32
37.59
36.96
36.14
36.36
35.54
35.98
29.27
29.36
34.97
43.02
37.45
37.04
40.55
60.92
51.63
0.47
0.51
0.48
0.42
0.41
0.46
0.47
0.46
0.13
0.16
3.04
0.80
1.81
1.66
1.48
2.02
2.82
0.65
1.36
0.75
0.90
0.97
0.93
0.69
0.53
0.10
0.21
2.68
0.57
14.15
10.24
13.86
47.11
9.45
2.47
3.77
3.07
3.81
3.24
4.32
3.62
5.67
3.64
4.91
24.87
2.90
14.63
11.23
10.39
0.81
6.75
39.68
39.66
36.13
38.60
37.79
34.69
36.70
35.63
1.48
9.17
198.02
31.91
33.50
31.97
229.37
76.94
30.25
0.02
0.03
0.04
0.04
0.04
0.04
0.02
0.01
0.00
0.00
0.00
0.01
0.14
0.20
0.03
0.05
0.01
18.29
16.04
8.63
9.68
9.13
11.87
22.44
34.20
44.59
129.89
236.07
65.76
3.67
9.17
2.97
0.74
14.90
1.36
1.35
1.29
1.20
1.25
1.22
1.14
1.90
4.12
2.55
0.89
0.64
0.93
0.93
1.26
0.83
1.12
0.72
0.55
0.67
0.62
0.62
0.62
0.68
0.82
0.63
0.82
0.17
0.82
0.23
0.25
0.27
0.14
0.47
0.70
0.72
0.66
0.63
0.69
0.64
0.64
1.06
0.07
0.41
0.23
1.10
1.77
1.46
2.20
3.40
3.66
0.77
1.31
0.41
0.74
1.01
1.73
2.23
3.59
5.45
7.21
298.2
4.88
58.85
14.41
83.59
56.09
114.7
Guryul Ravine Section
Period
Triassic
Fm.
Kunamuh
Bed
53
6F
6A
5X
5W
5V
5U
5T
5R
5Q
5P
5N
Height(m)
49.1
46.8
45.5
45
44.6
44.1
43.7
42.8
42.3
41.4
41
Detrital Supplay
Productivity Proxies
Paleore
Th/U
Cr/
Th
Zr/
Cr
Zr
REE + Y
Y/Ho
Zr/Hf
CuEF
NiEF
Ni/
Co
Ba/
Al(x10)-4
Mn/
Fe
Pb
V/
Cr
V/(V
+ Ni)
2.69
1.87
1.02
1.16
3.39
2.09
1.46
4.82
2.95
2.86
6.33
35.61
35.17
34.58
57.88
25.77
44.77
87.33
51.65
40.72
25.98
9.18
0.51
0.44
0.40
0.26
0.39
0.30
0.16
0.26
0.34
0.51
0.93
25.46
25.47
12.55
20.29
30.68
15.52
18.25
58.65
36.22
33.62
71.55
69.10
69.08
74.92
76.13
163.85
78.56
81.28
174.30
114.96
118.78
156.51
31.34
29.86
31.12
31.32
28.26
32.61
31.30
28.39
28.77
33.81
31.08
37.19
35.76
35.54
37.65
37.01
37.36
37.21
39.42
37.08
36.90
34.62
0.54
0.40
0.53
0.45
0.56
0.53
0.47
0.58
0.43
0.40
0.26
1.43
0.92
2.47
1.76
1.35
2.95
1.57
1.00
0.95
1.16
1.01
2.77
2.53
4.76
3.62
2.11
3.79
2.22
2.09
2.45
1.95
2.71
6.61
19.33
12.22
10.96
17.71
13.00
17.67
25.38
23.31
31.35
24.92
0.06
0.02
0.11
0.13
0.03
0.17
0.08
0.02
0.04
0.04
0.01
4.33
2.96
2.07
2.73
4.82
3.44
3.35
7.30
2.93
10.30
7.86
0.51
0.69
0.51
0.23
0.32
0.24
0.21
0.23
0.38
0.39
0.88
0.61
0.72
0.54
0.53
0.60
0.40
0.58
0.72
0.71
0.56
0.63
(Continued on
V-EF
MoEF
0.94
2.39
1.03
1.78
1.24
4.15
0.84
2.74
0.87
5.27
0.83
9.24
0.94
2.61
1.09
1.81
0.98
3.22
0.64
2.26
0.74
0.81
following page)
Hg Chemostratigraphy Permian-Triassic Boundary
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Sample
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TABLE 1 | (Continued) Trace element (ppm) and elemental ratios (including rare-earth elements) for the Meishan and Guryul Ravine sections.
Guryul Ravine Section
Period
Fm.
Permian
Bed
50
49
48
47
Zewan
46
5
45
44
43
Sample
40.5
39.6
39.2
38.7
38.3
37.8
37.4
36.9
36.5
36
35.6
35.1
34.7
34.2
32.9
32.4
31.5
31.1
30.6
29.3
28.8
27.9
27.5
25.7
Detrital Supplay
Productivity Proxies
Paleore
Th/U
Cr/
Th
Zr/
Cr
Zr
REE + Y
Y/Ho
Zr/Hf
CuEF
NiEF
Ni/
Co
Ba/
Al(x10)-4
Mn/
Fe
Pb
V/
Cr
V/(V
+ Ni)
V-EF
MoEF
4.02
2.87
7.08
4.36
7.51
6.06
7.24
16.12
5.34
6.02
11.14
8.50
8.83
13.32
6.63
7.12
7.73
14.59
3.51
16.68
6.45
12.92
5.84
11.17
27.68
23.47
12.72
11.50
9.66
12.13
9.34
17.52
17.21
19.77
14.51
18.42
17.72
21.08
16.83
9.33
9.87
7.54
15.36
25.53
13.98
22.55
14.52
15.21
0.48
0.65
0.84
0.82
0.90
0.84
0.89
0.37
0.39
0.26
0.28
0.23
0.28
0.20
0.28
0.46
0.42
0.46
0.32
0.12
0.30
0.19
0.32
0.29
40.96
50.71
178.54
50.37
151.75
174.22
147.46
17.50
13.29
30.41
65.37
35.38
48.93
59.76
30.90
44.47
44.19
68.13
29.69
24.86
22.77
64.70
29.08
47.69
115.57
113.37
302.13
176.58
245.47
260.52
190.58
56.37
78.19
177.72
145.58
189.24
179.86
142.54
169.85
175.59
160.91
227.95
149.78
89.20
123.92
126.24
137.83
168.05
33.60
31.48
27.70
32.44
26.36
26.41
26.94
31.18
31.99
31.37
25.74
29.62
30.10
27.35
30.84
33.41
31.51
28.27
30.65
30.91
34.04
28.82
32.96
25.32
37.45
36.87
35.98
36.67
34.59
37.07
37.26
44.90
38.48
35.22
34.23
36.08
36.42
37.92
35.49
36.66
36.12
34.48
37.48
33.04
34.63
34.48
34.66
33.61
0.52
0.76
0.45
0.22
0.32
0.40
0.31
0.33
0.51
0.15
0.20
0.23
0.22
0.18
0.22
0.23
0.14
0.14
0.17
0.14
0.16
0.17
0.25
0.20
1.43
1.44
0.60
1.08
0.74
0.74
0.52
6.24
3.81
0.93
0.76
0.93
0.85
0.79
0.95
1.02
0.82
0.64
1.18
0.64
1.08
0.70
1.06
0.95
2.41
1.69
1.90
2.33
2.66
2.35
1.93
4.19
5.03
2.54
1.95
2.30
2.44
1.90
2.28
2.21
3.10
2.40
3.44
3.12
4.01
2.81
2.82
3.23
19.91
15.18
45.67
19.62
38.35
41.43
56.08
20.49
21.52
17.25
18.78
23.68
20.60
22.11
20.47
40.87
16.80
19.61
25.13
27.57
39.81
35.45
67.81
29.41
0.03
0.02
0.00
0.02
0.00
0.00
0.00
0.08
0.07
0.10
0.01
0.07
0.08
0.02
0.08
0.06
0.07
0.03
0.06
0.05
0.07
0.01
0.06
0.01
3.02
3.31
10.65
5.58
11.02
20.29
27.23
5.86
5.23
6.80
20.76
13.69
10.46
18.70
8.79
9.76
7.90
16.06
6.42
10.30
7.89
17.35
10.68
15.04
0.48
0.66
0.69
0.88
0.73
0.73
0.97
2.21
1.75
0.50
0.62
0.36
0.33
0.31
0.44
0.75
0.55
0.55
0.36
0.22
0.43
0.26
0.53
0.56
0.58
0.58
0.77
0.61
0.71
0.76
0.84
0.85
0.81
0.81
0.82
0.76
0.80
0.81
0.76
0.82
0.81
0.81
0.69
0.77
0.71
0.82
0.76
0.74
0.83
0.86
0.86
0.72
0.79
1.02
1.12
15.70
7.18
1.68
1.54
1.26
1.44
1.44
1.30
1.95
1.47
1.18
1.14
0.91
1.11
1.40
1.41
1.14
2.08
1.05
0.87
1.04
0.83
1.11
1.08
10.64
6.57
4.07
2.82
4.57
6.10
6.58
3.70
3.23
3.54
2.45
5.96
6.30
3.78
7.64
4.59
2.76
June 2021 | Volume 9 | Article 651224
Hg Chemostratigraphy Permian-Triassic Boundary
5M
5L
5J
5I
5H
5G
5F
5E
5D
5C
5B
5A
4Z
4Y
4X
4U
4S
4R
4Q
4N
4M
4K
4J
4F
Height(m)
Sial et al.
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TABLE 2 | δ13Ccarb (‰), δ13Corg (‰), total Hg [(ng/g), S (%) and Al (%) concentrations; Hg isotopes (‰) (Meishan Section].
(A) Meishan Section
Fm.
Bed
Sample
Height
(m)
δ13Ccarb
(‰)
δ13Corg
(‰)
THg
(ng/g)
TOC
(%)
S
(%)
Al
(%)
Hg/TOC
(ng/g/%)
Hg/S (ng/
g/%)
Hg/Al (ng/
g/%)
δ202Hg
(‰)
∆199Hg
(‰)
∆201Hg
(‰)
Triassic
Yinkeng
30
29D
Meishan 41
Meishan 40
MS29+10cm*
MS29*
Meishan 38
MS28*
Meishan 36
MS27cd
(PTB)*
Meishan 33
Meishan 29
MS27ab*
Meishan 26
MS26e*
MS26d*
Meishan 24
MS26c*
MS26b*
MS26a*
MS25g*
MS25f*
Meishan 51B
MS25de*
Meishan 51A
MS25bc*
MS25a
(EPME)*
Meishan 50
MS24fo*
Meishan 23
MS24ec*
Meishan 20
MS24ea*
Meishan 15
MS24d*
68.7
56.5
49.0
39.0
36.5
26.2
25.5
20.7
—
0.88
—
—
1.2
—
0.1
—
−24.97
−27.64
—
—
−26.5
—
−25.68
—
26.8
32.3
38.6
27.3
6.5
60.1
9.3
20.6
0.05
0.12
0.2
0.86
0.05
—
0.05
0.14
1.55
1.07
—
—
0.61
—
0.76
—
4.84
3.94
—
—
3.23
—
2.72
—
536.00
269.17
193.00
31.74
130.00
—
186.00
147.14
17.26
30.27
—
—
10.58
—
12.29
—
5.54
8.20
—
—
2.01
—
3.42
—
—
−1.39
−0.58
−1.07
—
0.00
0.03
−0.02
—
−0.09
−0.04
0.01
−1.47
—
−1.1
−0.03
—
−0.07
−0.07
—
−0.06
20.5
15.0
13.7
10.5
8.7
7.1
6.5
5.9
4.9
3.4
2.9
2.1
1.9
0.9
0.80
0.3
0.0
0.7
0.1
—
0.0
—
—
−1
—
—
—
—
—
—
—
—
—
—
−26.22
−26.64
—
−26.23
—
—
−30.49
—
—
—
—
—
−26.26
—
−28.52
—
—
16.03
13.49
19.4
46.9
103.2
139
122.3
129.4
140.4
156.8
115.2
71.7
32.95
61.4
32.9
75.5
59.5
0.05
0.04
0.16
0.07
0.56
0.94
0.57
1.16
0.81
0.82
0.82
0.54
0.05
0.34
0.49
0.18
0.09
0.72
0.69
—
1.14
—
—
0.02
—
—
—
—
—
1.72
—
1.48
—
—
2.44
3.04
—
6.12
—
—
7.76
—
—
—
—
—
13.57
—
11.08
—
—
320.60
337.25
121.25
670.00
184.29
147.87
214.56
111.55
173.33
191.22
140.49
132.78
659.00
180.59
67.14
419.44
661.11
22.11
19.64
—
41.00
—
—
5541.46
—
—
—
—
—
19.19
—
22.31
—
—
6.56
4.43
—
7.66
—
—
15.77
—
—
—
—
—
2.43
—
2.97
—
—
—
−1.87
−0.93
−1.07
−1.14
−1.55
−1.1
−1.54
−1.36
−1.29
−1.41
−1.72
−1.97
−2.08
−1.38
−1.73
−1.95
—
−0.08
−0.08
−0.03
−0.04
−0.05
−0.08
−0.12
−0.05
−0.04
−0.03
0.02
0.05
−0.03
−0.06
−0.01
0.12
—
−0.05
−0.07
−0.07
−0.08
−0.08
−0.10
−0.02
−0.1
−0.02
−0.03
0.02
0.04
−0.03
−0.1
−0.04
0.08
−0.1
−0.8
−3.5
−5.8
−7.5
−9.8
−22.5
−40.9
—
—
1.9
—
2.2
—
2.1
—
—
—
−28.03
—
−27.82
—
−29.59
—
346.4
176
5.24
27.5
11.76
15.3
8.8
61.5
0.11
0.19
0.09
0.97
0.20
0.44
0.20
0.56
1.99
—
0.13
—
0.71
—
0.09
—
2.29
—
0.17
—
0.26
—
0.19
—
3149.09
926.32
58.22
28.35
58.80
34.77
44.00
109.82
174.42
—
39.46
—
16.45
—
102.28
—
151.17
—
30.04
—
45.10
—
47.01
—
−1.79
−1.67
−2.61
−1.34
−1.62
−1.14
—
−0.65
−0.02
0.09
0.09
0.09
0.09
0.03
—
0.09
−0.06
0.02
0.08
0.05
0.13
0.05
—
0.08
29C
27D
27C
27B
Permian
27A
26
6
26
Changxing
25
—
—
24E
24E
24E
24D
June 2021 | Volume 9 | Article 651224
(B) Guryul Ravine
Period
Fm.
Bed
Sample
Height
(m)
δ13Ccarb
(‰)
THg
(ng/g)
TOC
(%)
S (%)
Al (%)
Hg/TOC (ng/
g/%)
Hg/S (ng/
g/%)
Hg/Al (ng/
g/%)
δ202Hg
(‰)
Triassic
Kunamuh
53
6F
6A
5X
5W
5V
49.05
46.8
45.45
45
44.55
1.22
1.5
0.88
0.49
0.76
10.01
4.95
7.58
8.83
6.54
0.17
0.23
0.18
0.19
0.18
0.20
0.02
0.03
0.02
0.03
1.62
2.34
0.77
1.27
1.78
58.88
21.52
42.11
46.47
36.33
0.50
3.26
2.91
3.92
2.45
6.17
2.11
9.80
6.94
3.66
—
—
—
−2.11
—
∆199Hg
(‰)
∆201Hg
(‰)
—
—
—
—
—
—
0.03
0.00
—
—
(Continued on following page)
Hg Chemostratigraphy Permian-Triassic Boundary
Period
Sial et al.
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TABLE 2 | (Continued) δ13Ccarb (‰), δ13Corg (‰), total Hg [(ng/g), S (%) and Al (%) concentrations; Hg isotopes (‰) (Meishan Section].
(B) Guryul Ravine
Period
Fm.
Bed
52
51
Permian
50
49
48
47
7
Zewan
46
45
June 2021 | Volume 9 | Article 651224
Height
(m)
δ13Ccarb
(‰)
THg
(ng/g)
TOC
(%)
S (%)
Al (%)
Hg/TOC (ng/
g/%)
Hg/S (ng/
g/%)
Hg/Al (ng/
g/%)
δ202Hg
(‰)
∆199Hg
(‰)
∆201Hg
(‰)
5U
5T
5R
5Q
5P
5N
5M
5L
5J
5I
5H
5G
5F
5E
5D
5C
5B
5A
4Z
4Y
4X
4U
4S
4R
4Q
4N
4M
4K
4J
4F
44.1
43.65
42.75
42.3
41.4
40.95
40.5
39.6
39.15
38.7
38.25
37.8
37.35
36.9
36.45
36
35.55
35.1
34.65
34.2
32.85
32.4
31.5
31.05
30.6
29.25
28.8
27.9
27.45
25.65
0.98
1.59
0.52
−0.51
−3.99
−0.52
−1.19
−2.07
−2.41
−0.36
−3.48
0.92
−3.62
1.09
0.99
1.17
−0.47
0.72
1.62
−0.65
0.54
0.86
3.13
−0.63
0.66
0.03
2.03
0.95
2.25
0.32
6.96
16.35
4.82
3.79
8.74
4.3
4.19
3.14
11.51
3.68
7.27
11.06
9.79
1.48
0.74
2.38
4.52
3.79
2.28
5.05
2.67
3.19
1.94
3.39
2.21
2.92
2.99
4.06
4.54
5.87
0.23
0.18
0.21
0.91
0.22
0.27
0.24
0.19
0.37
0.20
0.46
0.55
0.23
1.03
0.16
0.87
0.77
0.24
0.22
0.74
0.21
1.03
0.86
0.17
0.17
0.16
0.20
0.24
0.23
0.86
0.02
0.03
0.02
0.02
0.19
0.02
0.13
0.45
0.00
0.08
0.01
0.02
0.00
0.05
0.05
0.06
0.01
0.02
0.01
0.01
0.04
0.01
0.02
0.01
0.01
0.01
0.01
0.01
0.04
0.01
0.88
1.58
2.83
2.47
2.46
5.55
2.91
3.56
10.24
4.56
9.37
8.99
8.67
0.40
0.49
2.08
5.68
2.63
2.42
3.94
2.26
2.23
2.34
4.11
1.79
2.91
1.73
3.65
2.06
4.80
30.26
90.83
22.95
4.16
39.73
15.93
17.46
16.53
31.11
18.40
15.80
20.11
42.57
1.44
4.63
2.74
5.87
15.79
10.36
6.82
12.71
3.10
2.26
19.94
13.00
18.25
14.95
16.92
19.74
6.83
3.43
6.14
3.00
2.38
0.47
1.77
0.31
0.07
32.10
0.47
13.90
6.90
92.45
0.31
0.15
0.39
6.07
2.34
1.92
4.73
0.64
4.76
1.08
5.22
3.35
2.07
2.22
4.66
1.21
5.12
7.94
10.37
1.70
1.54
3.56
0.77
1.44
0.88
1.12
0.81
0.78
1.23
1.13
3.67
1.50
1.15
0.80
1.44
0.94
1.28
1.18
1.43
0.83
0.83
1.24
1.00
1.73
1.11
2.21
1.22
—
−1.87
−2.66
—
−1.75
−1.56
−1.46
—
−1.44
−2.04
−1.97
—
−2.15
−1.88
—
—
−2.35
—
—
—
−1.53
—
−1.50
−3.46
—
—
—
—
—
—
—
0.02
0.01
—
0.10
−0.03
0.06
—
0.04
-0.05
−0.03
—
0.04
0.04
—
—
0.11
—
—
—
0.13
—
0.05
0.18
—
—
—
—
—
—
—
0.00
−0.05
—
−0.03
−0.01
0.08
—
−0.01
0.01
0.02
—
0.05
0.01
—
—
0.08
—
—
—
0.12
—
−0.04
0.09
—
—
—
—
—
—
Hg Chemostratigraphy Permian-Triassic Boundary
44
43
Sample
Sial et al.
Hg Chemostratigraphy Permian-Triassic Boundary
FIGURE 2 | Trace-element and elemental-ratio chemostratigraphy encompassing: Th/U, Cr/Th, Zr/Cr (Zr) in ppm, ƩREE + Y, Y/Ho, Al/Ca and (Al) in wt%, Cu-EF,
Ni-EF, Ni/Co, Ba/Al, Mn/Fe (Pb) in ppm, V/Cr, V/(V + Ni), V-EF, Mo-EF, U-EF, Mo/U, Ce/Ce*, Eu/Eu* and Pr/Pr* for the (A) Meishan D (GSSP) and (B) Guryul Ravine
sections.
the past may have occurred over sufficient time to allow
assimilation into the oceans, and acidification may not have
stressed ocean biota to the present extent.
Studies on LPME sites suggest that the mass extinction
occurred rather abruptly. Similar to the mass extinction
recorded at the Cretaceous–Paleogene (K/Pg) boundary, one
or multiple asteroid or comet impacts were invoked as
potential triggers of the LPME (Retallack et al., 1998). The
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Tethyan realm was the main focus of investigation of this
hypothesis which was based on the following observations: a)
iridium enrichment and occurrence of microspherules near the
Permian–Triassic Boundary (PTB) (Xu and Zheng, 1993), b)
shocked quartz (Retallack et al., 1998), c) fullerenes with trapped
helium and argon with isotope ratios identical to carbonaceous
chondrites (Becker et al., 2001), d) gigantic release of sulfur from
the mantle by a bolide impact at the end-Permian (Kaiho et al.,
8
June 2021 | Volume 9 | Article 651224
Sial et al.
Hg Chemostratigraphy Permian-Triassic Boundary
FIGURE 3 | (A–B) MoEF vs. UEF covariation for the Meishan D and Guryul Ravine sections (based on Tribovillard et al., 2012; Sosa-Montes et al., 2017), in which
MoEF [(Mo/Al)sample/(Mo/Al)PAAS] and UEF [(U/Al)sample/(U/Al)PAAS]. The PAAS composition used is from Taylor and McLennan (1985). The diagonal lines represent
multiples of the Mo/U ratio of present-day seawater. The general pattern of MoEF vs. UEF covariation in unrestricted marine trend for modern eastern tropical Pacific is
from Tribovillard et al. (2012), modified by Sosa-Montes et al. (2017) (C) Crossplot of Ce/Ce* vs. Pr/Pr* for the Meishan and Guryul Ravine sections (after Bau and
Dulski, 1996) (D) log(Ce/Ce*) vs. Nd (ppm) plot for both sections (modified from Wang et al., 2014).
2001, 2006), e) chondritic meteorite detritus (Basu et al., 2003),
and f) a supposed 250 Ma impact crater (Becker et al., 2004).
Altogether, this set of evidence has been questioned and
discredited to a certain point (e.g., Glikson, 2004; Koeberl
et al., 2004; Farley et al., 2005; Müller et al., 2005).
Finally, the P-Tr extinction entailed a substantial degradation
of marine benthic communities and extreme reduction of
sediment bioturbation, eradicating the sedimentary mixed
layer at an interregional scale (Hofmann et al., 2015). The
extinction of the infauna enhanced anoxia at the sedimentwater interface and was probably one of the causes of the
delayed recovery from the mass extinction during the Early
Triassic.
change from oxic to anoxic-euxinic conditions. High-resolution
δ 13C and Hg chemostratigraphic patterns from the
stratigraphically expanded, siliciclastic-dominated Guryul
Ravine succession were compared with the trends from the
Meishan GSSP limestone-dominated, condensed section.
Concentrations of major, trace, and Rare Earth Elements
(REE) and elemental ratios were used as paleoredox proxies
and to evaluate the detrital input. Hg stratigraphic patterns
were also contrasted with those from widely separated,
classical PTB sections and Hg isotope ratios were used to
unravel the source of the preserved Hg in the Guryul Ravine
and Meishan sections.
Aims of the Study
GEOLOGICAL SETTINGS
This study aims to contribute toward a better understanding of
the relationship between δ13C and Hg chemostratigraphy,
reinforcing the use of the latter as a reliable proxy in global
correlations, especially in the case of sedimentary sequences
deposited coevally with LIP volcanism. It also contributes to
the discussion on the normalization of Hg to Total Organic
Carbon (TOC), total S (TS as a surrogate of pyrite content),
or Al (i.e., surrogate of clay content) by providing valuable data
regarding situations when marine deposition witnessed, besides
coeval volcanism, a marine transgression accompanied by a redox
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Meishan Succession Global Stratotype
Section Point, China
The global stratotype section and point (GSSP) of the PTB has
been defined at the base of the Hindeodus parvus horizon of the
Meishan succession (base of bed 27c of the Meishan section) (Yin
et al., 2001). This succession in Changxing County, Zhejiang
Province, South China (Figure 1), has been subdivided into 115
beds (Li and Jones, 2017) and the PTB has an estimated
radiometric age of 251.902 ± 0.024 Ma, according to Burgess
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FIGURE 4 | Lithostratigraphy and conodont biozones for the Meishan D (GSSP) section, according to Chen et al. (2015); δ13Ccarb curve is from Chen et al. (2005)
and this study; δ13Corg and Hg/TOC chemostratigraphic pathways were modified from Sial et al. (2020a); Hg/S (Hg/Total sulfur, TS), Hg/Al and Δ199Hg‰ (MIF) variation
patterns are from this study. Gray stars along the Δ199Hg‰ (MIF) variation curve represent samples from Grasby et al. (2017). Variations in redox conditions are
according to Shen et al. (2007), based on the size distribution of framboidal pyrite.
et al. (2014). This condensed Tethyan carbonate-dominated
section in which the GSSP was defined renders correlation
between PTB sections difficult in high-resolution proxy
studies, as the extinction interval (Beds 25–28) is only 0.36 mthick (e. g., Baresel et al., 2017; Zuchuat et al., 2020).
At Meishan, the biodiversity crisis which marked the LPME is
marked just below Bed 25 by a significant loss of most fusulinids,
ammonoids and many brachiopods (Jin et al., 2000; Yin et al.,
2001; Erwin, 2006; Yin et al., 2007). The onset of biodiversity
reduction started during the deposition of bed 22 (Jin et al., 2000;
Yin et al., 2001; Yin et al., 2007) and coincides with a negative shift
in conodonts δ44/40Ca values in this section. This shift was also
observed in marine carbonates from several P–Tr sections
elsewhere (Payne et al., 2010; Silva-Tamayo et al., 2018) and
may reflect a major perturbation of the marine Ca cycle caused by
ocean acidification (Hinojosa et al., 2012). The perturbation of the
Ca isotope composition of seawater was probably a consequence
of massive emission of volcanic CO2 into the P–Tr atmosphere by
LIPs (e.g., Silva-Tamayo et al., 2018). The negative shift in
conodont δ 44/40Ca values was recorded up to the LPME (beds
25–26), and an additional one at the earliest Triassic Bed 28
concomitant with the ETME (Xie et al., 2005). A third ocean
acidification event probably postdated the LPME by 60 kyr (SilvaTamayo et al., 2018). These negative shifts of δ44/40Ca coincide
with a major perturbation of the marine U isotope cycle, probably
related to volcanism, ocean acidification and anoxia (Lau et al.,
2016).
The extinction interval at the P–Tr transition may contain
depositional hiatuses at the PTB or below (Cao and Zheng, 2007;
2009; Zheng et al., 2013; Shen et al., 2018), and corresponds to a
maximum of 61 ± 48 kyr according to Burgess and Bowring
(2015). The STLIP lava flows, sills, and explosively erupted rocks
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yielded U-Pb ages (Burgess and Bowring, 2015) that, coupled
with biostratigraphic evidence, point to a single, brief catastrophic
LPME event, confirming previous suggestions (Jin et al., 2000;
Rampino et al., 2000; Shen et al., 2011; Wang et al., 2014). These
results suggest that about 65% of the total lava/pyroclastic volume
was erupted in about 300 kyr before and concurrent with the
LPME, and eruptions continued for at least 500 kyr after the mass
extinction ceased. Voluminous STLIP sills intruded into a shallow
crust with huge amounts of carbon-rich sediments, probably
inducing the most important killing mechanism in the LPME
(stage 2, extrusive hiatus; Burgess et al., 2017).
Sulfur content and isotopic composition in pyrite near the
PTB in the Meishan section have been determined by Jiang et al.
(2006). Shen et al. (2007) were the first to report abundant pyrite
framboids from the upper part of bed 24 and from beds 25, 26,
and 29 (they are absent in bed 27), implying a dramatic decline in
benthic oxygen levels during the LPME. Framboids from the
Meishan succession display a narrow size distribution, with
average diameters from 4.6 to 8.7 μm, typical of pyrite
framboids formed under anoxic/dysoxic conditions. According
to these authors, based on the abundance and size of these
framboids, the redox conditions of deposition changed from
upper dysoxia (bed 24e, uppermost Changhsing Formation)
and lower dysoxia (pyrite lamina, beds 25 and 26, Yinkeng
Formation) to oxygenation (bed 27), and again to upper
dysoxia (Bed 28) and lower dysoxia (bed 29). A later study
reported negative δ34S values and negative Δ33S in pyrites
from Meishan, hinting at an origin from sulfate-deficient
waters, probably related to a near-shutdown of bioturbation,
due to shoaling of anoxic water during the LPME (Shen et al.,
2011). Recently, Wei et al. (2020) reported two gradual, negative
δ 13Ccarb shifts before and after the PTB. These authors suggest
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FIGURE 5 | Lithostratigraphy and conodont biozones for the Guryul Ravine section according to Brookfield et al. (2019); δ13Corg curve from Algeo et al. (2007);
δ 13Ccarb pathway is from Sial et al. (2020b), Δ199Hg‰ (MIF), Hg/TOC, Hg/S (Hg/TS), Hg/Al chemostratigraphic curves are from this study; variations in the redox
conditions are based on Huang Y. et al. (2019), two-marine anoxic events based on the distribution of framboidal pyrite. Relative depth of deposition (shallow, deep) is
from Singh et al. (2015).
that the Late Permian ocean redox conditions varied through
three successive oxic–dysoxic, dysoxic–euxinic, and highfrequency euxinic stages. Oxic and high-frequency euxinic
events occurred in the early Griesbachian and middle
Griesbachian, respectively. The short-term euxinic events were
not associated with the negative δ13Ccarb excursions. The secular
trends of aromatic hydrocarbons diagnostic of anoxygenic green
sulfur bacteria point to periods when euxinic conditions extended
into the photic zone during the entire Changhsingian stage (Cao
et al., 2009).
biostratigraphical understanding of these two sections. These
Tethyan sedimentary successions in Kashmir have experienced
greenschist-facies metamorphism (Herren, 1987; Dèzes, 1999;
Algeo et al., 2007).
The deposition of the Guryul Ravine succession occurred at
moderate sedimentation rates (∼10–20 m Myr−1: Algeo et al.,
2007) in an outer-shelf or deep-ramp setting in the Neotethys
(Brookfield et al., 2003; Wignall et al., 2005; Algeo et al., 2007;
Baud and Bhat, 2014). About 15 m-thick siliceous shales of the
Gangamopteris-beds cover flood basalts of the 289 Ma-old Panjal
Traps (Shellnutt et al., 2011, 2014) and are covered, in turn, by the
100 m-thick Zewan Formation. Based on the variation in
carbonate content, this formation was subdivided into four
members (A–D) and the overlying Khunamuh Formation, into
six members (E–J) according to Nakazawa et al. (1975). The
Guryul Ravine section exhibits gradual faunal changes across the
PTB, similar to the other Neotethys P–Tr sections (e.g., Nakazawa
et al., 1970; Teichert et al., 1970; Nakazawa and Kapoor, 1981;
Sheng et al., 1984; Brookfield et al., 2013, 2019; Baud and Bhat,
2014). Conodont biostratigraphy of this section can be found
in Murata (1981), Matsuda (1981, 1982, 1983, 1984), and
Wignall et al. (1996). Tewari et al. (2015) found evidence of
depauperate pollen and spore assemblages in the uppermost
Zewan Formation and a rich palynoflora in the basal
Khunamuh Formation. Whereas in the GSSP Meishan
section the LPME and the PTB horizons are separated by
0.36 m, in the Guryul Ravine there is a stratigraphically
expanded P–Tr transition with the two horizons separated
by ∼2.60 m (Algeo et al., 2007) or 3.10 m (Shen et al., 2011;
Brookfield et al., 2013). Therefore, the study of different
events within the P–Tr transition is considerably easier at
Guryul Ravine section. The LPME horizon encompasses the
top of the Zewan Formation (bed 46) and the base of the
Guryul Ravine Succession, India
The PTB Guryul Ravine section near Srinagar in the Kashmir
region in north India (Figure 1) was a candidate for the GSSP
before the selection of Meishan D (e.g., Kapoor, 1996; Yin et al.,
2001). Kashmir was once a part of the Gondwana supercontinent
at the southern side of the Paleotethys Ocean and adjacent to
Oman during the late Paleozoic (Brookfield et al., 2013).
Widespread marine transgression in northern India (e.g.,
Kapoor, 1992; Garzanti et al., 1998) followed the rapid
subsidence of the northern Gondwana margin with separation
of tectonic blocks and opening of the Neotethys Ocean (e.g.,
Brookfield, 1993; Garzanti et al., 1996; Brookfield et al., 2013,
2019). The Guryul Ravine section and the Spiti Valley, another
remnant of the peri-Gondwanan shelf in India (Ghosh et al.,
2016), are probably the two best geologically documented
Neotethys Ocean PTB sections in India. Geological and
paleoredox aspects of the Guryul Ravine and the Spiti sections
are well documented (e.g., Murata, 1981; Matsuda, 1981, 1982,
1983, 1984; Brookfield et al., 2003, 2013, 2019; Wignall et al.,
2005; Algeo et al., 2007; Korte et al., 2010; Tewari et al., 2015;
Ghosh et al., 2016; Kumar et al., 2017; Huang Y. et al., 2019),
resulting in significant advances in the sedimentological and
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Khunamuh Formation (bed 47). The PTB, in turn, is recorded
at the top of bed 51 and the base of bed 52, within the
Khunamuh Formation (Algeo et al., 2007).
regarded as key sections for unraveling the redox history of the
Neothethys Ocean across the PTB.
Reliable biostratigraphic correlations have been established
between the Guryul Ravine and Meishan sections (e.g., Chen
et al., 2005; Algeo et al., 2007; Brookfield et al., 2013). However,
seismites and tsunamites present at the Guryul Ravine seem to be
absent in the Meishan GSSP. Identical structures have not been
observed in any other PTB section in India, except for
Mandakpal. Those described at Guryul Ravine have been
rejected as such by Krystyn et al. (2014).
Correlation Between Guryul Ravine and
Meishan Sections
Correlations between biostratigraphic patterns of the Guryul
Ravine and Meishan sections were reported previously (Chen
et al., 2005; Algeo et al., 2007; Brookfield et al., 2013). The Guryul
Ravine beds D43–D46-1 correlate with bed 24 at the Meishan
section, and Guryul Ravine beds D46-3–E51 match beds 25–27 at
Meishan (Chen et al., 2005). The Guryul Ravine section displays,
similar to the Meishan section, an extinction pattern with the
largest extinction rate between the LPME and the PTB and a less
intense rate at the ETME (Algeo et al., 2007). The main extinction
event in this section is recorded in Unit E1 (bed 47) and the less
intense ETME at the base of Unit E2 (bed 52; Shen et al., 2006). A
continuous sea-level rise across the LPME, i.e., from the base of
bed 46 up to the base of bed 52 (PTB horizon), was reported by
Algeo et al. (2007). This transgression was caused either by a
eustatic sea-level rise or active vertical tectonics at the northern
Gondwanan margin (Hallam and Wignall, 1999). Its maximum
flooding seems to have been characterized by an anoxic event
(Shen et al., 2006).
Geochemical paleoredox proxies (i.e., Mo/U, V/Cr, Mn/Fe,
Ce/Ce*, Eu/Eu*) suggest that the environment in the Guryul
Ravine section was oxic or suboxic across the PTB, but marine
anoxia with increasing clay content was probably developed
during the LPME event, as documented in several
investigations (e.g., Algeo et al., 2007; Brookfield et al., 2013).
Wignall et al. (2005) observed a change from non-laminated,
pyrite-poor mudstones to laminated, pyrite framboid-rich shales
in bed E49 of the Khunamuh Formation, 1 m above the LPME
horizon. Additionally, the TOC to total phosphorus molar ratio
(TOC/P) is < 10 in the late Permian and >10 in the early Triassic,
suggesting a transition from suboxic to anoxic conditions at the
PTB (Kumar et al., 2017).
Two pronounced stages of oceanic oxygen deficiency at the
Guryul Ravine section were documented by pyrite framboid size
and morphology (Huang Y. et al., 2019): i) during the latest
Permian Hindeodus praeparvus–Clarkina meishanensis biozone
and ii) during the earliest Triassic Isarcicella staeschei biozone.
Huang Q. et al. (2019) proposed that these two anoxic events
correlate with anoxic events at the Meishan GSSP and should be
considered as global characteristics of the P–Tr transition.
Wignall et al. (2005) recognized a possible connection between
the LPME and deep-water anoxia at the Guryul Ravine section.
Besides, framboidal pyrite and phosphatic nodules in the Late
Permian shales of the Gungri Formation in the Spiti Valley
(India) also suggest anoxic to euxinic bottom water conditions
(Singh, 2012; Ghosh et al., 2016). These paleoredox conditions
contrast with those observed in other peri-Gondwanan sections
in which shallow-water anoxia took place at the ETME. The
Guryul Ravine and the Spiti Valley PTB sections are the only ones
in the Neothethys Ocean that record deep-water anoxia at the
P–Tr transition (Wignall et al., 2005). Therefore, these can be
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MATERIAL AND METHODS
Samples
The Meishan D section is located in the Meishan quarry, Zhejiang
Province, SE China, with coordinates 31°4′55″ N and
119°42′22.9″ E (see Jin et al., 2006). From the Meishan D
section, 33 samples were analyzed for C and Hg isotopes,
organic C, S, Al, and Hg concentrations. Out of these, 19
samples were also analyzed for REE + Y and redox-sensitive
trace elements.
The samples (n 109) from the Guryul Ravine section were
collected along a traverse perpendicular to the strata of the Zewan
Formation, starting at about 71 m above basalts of the Panjal
Traps, at the southeast side of the Guryul Ravine (coordinates of
the start point: 34° 04′ 24.13″ N and 74° 56′ 43.42″ E, and end
point: 34° 04′ 6.66″ N and 74° 56’ 48.20” E).
Analytical Methods
Major, Trace and Rare-Earth Elements
Major, trace and Rare-Earth Element (ME, TE, REE)
concentrations were determined in solution using inductively
coupled plasma mass spectrometer (ICP-MS) at the Geological
Survey of Denmark and Greenland (GEUS) in Copenhagen,
following the protocol of Holland et al. (2003). Aliquots of
pulverized samples were treated with HF and HNO3 in a
PTFE (Savillex) bottle at 130°C. A sample of the resulting
solution was dried on a hotplate (100°C, 24 h). The residue
was dissolved twice in HNO3 and dried (100°C, 24 h). Then
HNO3, internal standard solution (containing Ge, Rh, Re), and
distilled water (DIW) were added and allowed to react in a closed
PTFE bottle (130°C, 12 h). The resulting solution was diluted (10
× DIW to 50 ml) and analyzed for TE and REE using a Perkin
Elmer Elan 6100 DRC ICP-MS instrument. Instrumental
calibration was performed using solutions containing certified
concentrations of REEs and other elements. The reproducibility
of the measurement of the element concentrations, assessed by
two relative standards deviations (2RSD), were typically lower
than 3% as determined from multiple analyses (n 4) of in-house
standards Disko-1 and international reference materials BHVO-2
and BCR-2. The elements Ni, Mo, Cs, and Eu have higher
analytical uncertainties, but were consistently determined with
2RSD better than 10%. Detection limits of the method are listed
and were on average six times lower than the lowest
concentration measured in any sample. For some elements, in
particular Sc, Zn, Ga, Nb, Eu, Ho, Tm, Yb, Lu, Hf, Ta, Pb, and Th,
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the detection limit was relatively higher (5% percentile of
measured values in the samples).
present study. For the Guryul Ravine section, Hg concentration
analyses for 35 samples were performed at LABOMAR (Institute
of Marine Sciences, Fortaleza, Brazil) following the protocol
detailed previously (Sial et al., 2020a). Shortly, Mercury
concentrations were determined in well homogenized
powdered samples. To avoid cross-contamination and ensure
low blanks, all the glass and plastic labware were decontaminated
(24 h immersion in 10% v/v Extran solution of Merck; 24 h
immersion in 5% v/HCl solution; final rinse with purified
water using a Milli-Q system). The chemical reagents were of
analytical grade purity or higher (Moynier et al., 2020). At
LABOMAR, the Mercury extracted (BrCl/HCl oxidant/acid
solution) as Hg2+ was reduced to volatile Hg0 with SnCl2 and
detected by cold vapor atomic fluorescence spectrometry
(Millennium Merlin PSA instrument). The analyses were done
in duplicates, and the results reproducibility was <10%. The Hg
determination accuracy was assessed by replicate analysis of
reference material NIST 1646a (Sigma Aldrich); the
reproducibility and accuracy were <5%. The limit of Hg
detection (LOD) was 0.02 ng g−1 Hg and the limit of
quantification was 0.06 ng g−1 Hg. Blank signals were lower
than 0.5% of the average sample signal.
At Tianjin University, the Mercury concentrations were
determined using a Lumex Hg analyzer RA 915F equipped
with a pyrolysis attachment (PYRO-915+). The reference
materials CRM GBW07311 (stream sediments) and
GBW07405 (yellow red-soil) from the National Center for
Standard Materials (Beijing, China) were analyzed for every
five unknown samples for quality control, with the determined Hg
concentration within ±10% of the certified values. The instrumental
LOD was of 0.5 ng g−1 Hg. The dual-stage combustion system and
trapping method described by Huang et al. (2015) have been adopted
for purification and pre-concentration of Hg. Briefly, each weighed
sample was step-combusted in a quartz tube in the first combustion
furnace, then the combustion products (Hg0) were carried by Hg-free
O2 gas through the second decomposition furnace of temperature set
at 950°C, finally trapped by a 3 ml 40% (v/v, 2HNO3/1HCl) acid
solution (Huang et al., 2015). This method was validated with
reference materials (CRM) GBW07311 (Hg 72 ± 9 ng/g) and
GBW07405 (Hg 290 ± 30 ng/g), which yielded recoveries of
98 ± 5% (2SD, n 8) and 105 ± 7% (2SD, n 20), respectively.
The detectable Hg in the procedural blank (<0.05 ng, n 2) of this
protocol method was negligible compared to the amount of total Hg
(>3 ng) in samples. To assess eventual systematic differences in Hg
concentrations between analytical procedures, a few samples were
analyzed at both labotaories (Tianjin and LABOMR), and the
obtained Hg concentration differences were within the analytical
uncertainty.
Total Carbon and Total Sufur
Total carbon (TC) and total sulfur (TS) analyses were performed
on whole-rock powdered samples using a LECO CS230 induction
furnace at GEUS, Copenhagen. The analytical reproducibility was
better than ±0.5% for both TC and TS.
Total Organic Matter
Organic matter analyses were performed using a HAWK
pyrolysis equipment (a Rock-Eval 6 equivalent instrument
from Wildcat instruments and services, Humble, TX) at
GEUS, Copenhagen. Calibration was done with the IFP
150,000 standard and in-house standard. The analytical error
for the total organic carbon (TOC, wt%) measurements with a
HAWK instrument is similar to that obtained by Rock-Eval 6
pyrolysis and oxidation, and generally lower than 2%.
Organic Carbon Isotopes
The organic carbon isotope ratios (δ13Corg, ‰ VPDB) were
determined on decarbonated (10% HCl, 60°C) samples by
elemental analysis and isotope ratio mass spectrometry (EA/
IRMS), using an IsoPrime triple collector isotope ratio mass
spectrometer at the Department of Geosciences and Natural
Resource Management, University of Copenhagen. The carbon
stable isotope ratios were reported in the delta (δ) notation as the
per mil (‰) deviation relative to the Vienna Pee Dee Belemnite
standard (VPDB). The measured δ 13C values were normalized to
the VPDB scale using the in-house AKsill-9 standard (δ 13C:
−25.30‰). The reproducibility of the δ13Corg values was better
than 0.1‰.
Carbonate Carbon and Oxygen Isotopes
The carbonate carbon and oxygen isotope ratios (δ 13Ccarb and
δ 18Ocarb ‰ VPDB) of whole-rock samples from the Guryul
Ravine section were determined at the Stable Isotope
Laboratory (LABISE) at the Department of Geology, Federal
University of Pernambuco, Brazil. Extraction of CO2 gas was
performed using a conventional high vacuum extraction line after
reaction with 100% orthophosphoric acid (25°C, 24 h). The CO2
gas recovered from the samples was analyzed in a Delta V
Advantage isotope ratio mass spectrometer (Thermo Fisher
Scientific, Bremen, Germany). The measured δ 13Ccarb values
were normalized by calibrating with the reference CO2 gas
with international carbonate isotope reference materials (NBS18 and NBS-19). Analytical uncertainty (1 sigma) monitored by
replicate analyses of NBS-19 and the in-house laboratory
standard BSC (Borborema skarn calcite) was not greater than
±0.05‰ for δ 13Ccarb and ±0.1‰ for δ18Ocarb.
Mercury Isotopes
The Hg isotope compositions were determined using a multicollector ICP-MS (MC-ICP-MS, Nu Plasma 3D, Nu Instruments,
United Kingdom) equipped with a home-made continuous flow
cold vapor generation system at the Institute of Surface-Earth
System Science, Tianjin University, following the same setup as
described in previous studies (Chen et al., 2010; Huang et al.,
2015; Zhang et al., 2020). Briefly, the Hg(0) vapor was generated
Mercury Concentration
In this study, 15 new analyses of Hg concentration for the
Meishan section were performed at the laboratories of the
School of Earth System Science of Tianjin University (China).
Additionally, 19 analyses of Hg concentration in samples from
this section, available in Grasby et al. (2017) were also used for the
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FIGURE 6 | A NW–SE Hg/TOC transect across the Pangea supercontinent (252 Ma ago) encompassing nine PTB sections, from Ursula Creek (Canada) through
Hovea-3 (Australia), including Meishan D and Guryul sections based on Sial et al. (2020a). In building this transect, the position of the LPME and ETME horizons have
been kept as shown in their original Figures and the PTB datum was used for correlation.
ΔX Hg δX Hg%° − βX × δ 202 Hg
via online reduction by SnCl2 (3%) solution. The Tl aerosol was
produced from Aridus II nebulizer and simultaneously
introduced into the plasma with Hg(0). The instrumental mass
bias was corrected by an internal NIST-997 Tl standard and NIST3133 Hg standard-sample bracketing method. The samples were
introduced at 0.75 ml/min, which gave an instrumental sensitivity
of 2.8 V on 202Hg for 1 ng mL−1 Hg solution. Hg concentrations of
bracketing standards NIST-3133 and secondary standard UMAlmadén were matched to the sample solutions within 10%. All
samples were analyzed in three blocks with 33 cycles for each and with
10 min washout between samples to ensure that the blank levels were
<2‰ of the preceding sample signals.
The mass-dependent fractionation (MDF) of Hg isotopes is
represented as delta (δ) notation, the per mil deviation of the
x
Hg/ 198 Hg ratio relative to the NIST-3133 standard, and is
defined as:
where the mass-dependent scaling factor βx is 0.252, 0.502, and
0.752 for 199Hg, 200Hg, and 201Hg, respectively.
For quality assurance and control, replicate analyses of the UMAlmadén Hg standard yielded average values of −0.55 ± 0.09‰,
−0.01 ± 0.07‰, 0.02 ± 0.07‰, and −0.03 ± 0.10‰ for δ 202Hg,
Δ199Hg, Δ200Hg, and Δ201Hg, and −1.75 ± 0.05‰, -0.35 ± 0.03‰,
0.00 ± 0.05‰, and −0.32 ± 0.04‰ for those of CRM GBW07405 (2
SD, n 5), respectively. These values are consistent with previous
results (Chen et al., 2010; Huang Q. et al., 2019). The two SD values of
the replicate UM-Almadén δ202Hg, Δ199Hg, Δ200Hg, and Δ201Hg
determinations represent the analytical uncertainties, which were 0.08,
0.03, 0.03, and 0.05‰, respectively. Errors are reported as two SD
from the replicate (generally duplicate) measurements of the same
sample solution.
δ X Hg X Hg198 Hgsample X Hg198 HgNIST−3133 − 1 × 1000
RESULTS AND DISCUSSION
(1)
Trace Elements Chemostratigraphy
where x refers to the mass of each isotope between 199 and
202 amu.
Hg isotopes can also undergo mass-independent fractionation
(MIF), which are defined as the deviation of a measured delta
value from the theoretical MDF value, according to the
equation:
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(2)
A safe interpretation of the trace element and REE concentration
patterns and derived paleoproxies (i.e., Ce/Ce*, Eu/Eu*) would
require the knowledge of the hosting phases, and the diagenetic
history of the studied sediments and sedimentary rocks. This is
because these materials may have undergone multiple stages and
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FIGURE 7 | δ202Hg ‰ (MDF)–Δ201Hg (MIF) crossplot for (A) Meishan D section (analyses are partially from this study and partially from Grasby et al., 2015a), and
(B) Guryul Ravine section (this study). See text for explanation.
FIGURE 8 | (A) The Δ199Hg‰ (MIF) vs. (B) Hg (ng.g−1) crossplot for the two studied sections. See text for explanation.
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styles of post-depositional diagenetic (and probably also
hydrothermal)
transformations,
including
dissolution,
recrystallization, neomorphism, and phase changes during
which the (local) redox conditions and fluid compositions may
have varied widely from the overlying seawater (e.g., Tribovillard
et al., 2006; Hood et al., 2018; Liu et al., 2019; Algeo and Liu,
2020). Having this in mind, we present in Table 1 the abundances
of selected trace element (TE) of 19 samples from the Meishan D
section and 35 samples from the Guryul Ravine section. The TE
ratios used as paleoproxies are shown in Figure 2, along with REE
ratios (Ce/Ce*, Eu/Eu*, Pr/Pr*) and ΣREE + Y (Table 1). The
element enrichment factors (EFs) relative to PAAS were used to
compare the concentrations of authigenic V, Mo, and U in both
sections (Algeo and Tribovillard, 2009). A detectable authigenic
enrichment of element X corresponds to XEF > 3, and a
substantial one, to XEF > 10. The Mo and U concentrations
are shown as MoEF and UEF profiles in Figure 2 and MoEF vs. UEF
plots in Figures 3A,B. The seawater Mo/U molar ratio is shown
to guide interpreting relative enrichments of authigenic Mo vs.
authigenic U. The MoEF and UEF values {i.e., MoEF [(Mo/
Al)sample/(Mo/Al)PAAS] and UEF [(U/Al)sample/(U/Al)PAAS]} are
based on Tribovillard et al. (2012) and Sosa-Montes et al. (2017),
using the PAAS composition from Taylor and McLennan (1985).
The MoEF and UEF covariations have been used as indicators of
paleoredox conditions in marine systems (e.g., Algeo and Lyons,
2006; Tribovillard et al., 2006; Algeo and Tribovillard, 2009;
Algeo and Rowe, 2012; Tribovillard et al., 2012; Zhou et al.,
2012; Azrieli-Tal et al., 2014; Algeo and Liu, 2020). Algeo and
Tribovillard (2009) recognized multiple controls in the
MoEF−UEF covariation, mainly i) the concomitant increase of
sedimentary Mo and U accumulation with decreasing benthic
dissolved oxygen concentration in open-ocean upwelling regions
and ii) preferential uptake of Mo over U in weakly restricted
basins with a particulate shuttle in the water column. Therefore,
as the accumulation of Mo and U have conjoint hydrographic and
redox controls, the use of their concentrations and derived
proxies (i.e., Mo/U ratios, MoEF−UEF covariations) as a
paleoredox proxy needs to be supported by (Eq. 1)
comparison with other paleoredox proxies based on redoxsensitive/sulfide-forming elements (e.g., V, Ni, Cu, Zn) and
(Eq. 2) consideration of the hydrographic settings of the
studied paleoenvironment (Algeo and Lyons, 2006;
Tribovillard et al., 2006; Tribovillard et al., 2012). These
authors recognized the difficulty of such assessments (e.g.,
basin configuration, circulation patterns) due to the
incomplete knowledge of the paleogeographic dimensions of
many anoxic paleoenvironments.
The Meishan and Guryul Ravine sections display distinct Mo/
U and MoEF−UEF covariation trends across the PTB transition
(Figure 3). The diagonal lines in Figures 3A,B represent
multiples (between 0.1 and 3.0) of the Mo/U ratio of presentday seawater (SW) and the general pattern of MoEF−UEF
covariation in unrestricted marine trend for modern eastern
tropical Pacific (Tribovillard et al., 2012, modified by; SosaMontes et al., 2017). Two trends result from this plot: a) a
number of pre-LPME, extinction interval and post-ETME
samples from the Meishan section (Figure 3A) plot roughly
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along the 0.1 × (Mo/U) present-day SW line; b) samples from
the PTB and ETME horizons (suboxic) and three samples from
the extinction interval (suboxic-anoxic) plot along the 0.3 × (Mo/
U) of the present-day SW line and the LPME (euxinic), on the 3.0
× (Mo/U) SW diagonal line. The second trend follows
approximately the covariation in unrestricted marine trend for
modern eastern tropical Pacific. In the Guryul Ravine section
(Figure 3B), two trends are observed: a) most samples plot along
the 0.3 × (Mo/U) SW line (suboxic to anoxic) and b) four preLPME samples plot along the 1.0 × (Mo/U) SW line. The samples
from the LPME horizon at Guryul Ravine and at Meishan exhibit
contrasting behavior, oxic and anoxic, respectively.
At Meishan, Mo/U, V/Cr, Ce/Ce*, Eu/Eu*, and Pr/Pr* ratios
and Pb concentration show little variation in the Changhsing
Formation (bed 24), while VEF, MoEF, UEF, andV/(V + Ni) exhibit
strong fluctuations (Figure 2A; stratigraphic location of beds in
Meishan and Guryul Ravine sections are shown in Figures 4, 5).
All paleoredox ratios from this section (including Mo/U) display
a peak in the LPME horizon, followed upsection by a relatively
discrete variation in the Yinkeng Formation, except for the Ce/
Ce*, which increases towards the PTB horizon, and the Eu/Eu*
and Pr/Pr* ratios that undergo slight decrease (beds 26 and 27).
The slightly negative Eu anomaly in the LPME samples could
suggest precipitation from either fluids with lower temperature
than pre-LPME horizons, or chemical complexation reactions
which extended the stability of Eu3+ to lower oxygen
concentrations, preventing the reduction to Eu2+, and
therefore decreasing the Eu/Eu* values (e.g., Bau, 1991; Hood
et al., 2018).
In general, all of these paloredox ratios tend to show little
systematic variation towards the ETME horizon. The V/(V + Ni)
values, used as a redox paleoproxy (Arthur and Sageman, 1994),
are uniformly high (>0.60 on average) and suggest deposition
under anoxic conditions in the extinction interval (Figure 2).
Mn/Fe ratios >0.1 in the pre-LPME Bed 24 (Figure 2A) favor oxic
conditions, while values <0.1 at the top of Bed 24 and in beds 25
and 26 across the LPME suggest low redox conditions, according
to Arthur and Sageman (1994). At Meishan, the detrital proxies
Th/U, Zr/Cr, Zr, Al, Al/Ca, and also ΣREE + Y display a marked
positive excursion defined by several samples at the LPME
horizon (Figure 2). Apart from this, the detrital supply
proxies (i.e., Th/U, Zr/Cr, Al/Ca) show a little systematic
trend in the Changhsing Formation, while Cr/Th and Y/Ho
exhibit more pronounced variation. Some detrital input
proxies show a strong negative (Th/U, Z, and ΣREE + Y) or
positive (Y/Ho) excursion in the LPME–PTB interval in the
Yinkeng Formation (beds 26 and 27) and little variation from
the ETME upsection. The low Y/Ho ratios suggest a diagenetic
lithogenous geochemical signature (e.g., Zhang et al., 2016). The
productivity proxies Cu-EF, Ni/Co, and Ba/Al show strong
variation in the Changhsing Formation, with a positive peak at
the LPME horizon and slight variation within the extinction
interval (Figure 2A).
At Guryul Ravine, the element ratios show more pronounced
fluctuations than in the Meishan section (Figure 2B). Except for
VEF, all paleoredox proxies show strong pre-LPME changes (bed
46, Zewan Formation). Upsection, they follow a pattern similar to
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Meishan, punctuated by a marked increase in most redoxsensitive element ratios at the LPME horizon, except for Ce/
Ce*, which shows a slight decrease. Most paleoredox ratios [V/(V
+ Ni), MoEF, UEF, Ce/Ce*, Eu/Eu*] and Pb are enriched within the
PTB–ETME interval and the Pr/Pr* ratio exhibits more erratic
fluctuations toward lower values. The Mn/Fe > 0.1 in Bed 46
suggests pre-LPME oxic conditions, while Mn/Fe < 0.1 in bed 49
points to reducing conditions (Figure 2B). The Mn/Fe > 0.1 in
beds 51, 52, and at the base of Bed 53 suggest a return to oxic
conditions upsection, and the Mn/Fe < 0.1 values before the
ETME to more reducing conditions. The detrital input proxies
Th/U, Zr/Cr, Zr, and ΣREE + Y show some fluctuations in the
pre-LPME interval (bed 46) and a strong positive shift at the
LPME horizon. The increases in Zr and ΣREE + Y are in line with
Algeo et al. (2007). The productivity proxies (NiEF, Ni/Co and Ba/
Al) show strong variations and a positive peak at the LPME,
except for the CuEF with a discrete negative shift.
The ΣREE + Y concentrations vary from 10 ppm (Bed 24 D) to
143 ppm (Bed 24E) in the analyzed samples from the Changhsing
Formation and from 121 ppm (Bed 29C) to 430 ppm (Bed 26) in
samples from the Yinkeng Formation in Meishan (Table 1). In
Guryul Ravine, ΣREE + Y varies from 56 to 260 ppm (Bed 46) in
the Zewan Formation and from 78 (Bed 53) to 302 ppm (Bed 48)
in the Kunamuh Formation (Table 1). Sea-level oscillations
across the LPME horizon in the Guryul Ravine section have
probably allowed an important continental detrital influx and
subsequent ΣREE + Y enrichment. The strong co-variation
between Zr and ΣREE + Y in both sections (Figure 2)
suggests the presence of heavy minerals, especially zircon (e.g.,
Hoskin and Ireland, 2000; Bouch et al., 2002). The REE- and
likely Zr-enriched detrital phases were responsible for the REE in
these two sedimentary successions.
The Ce/Ce* values can shed some light on post-depositional
mobility of REE, as positive Ce/Ce* values indicate reducing
conditions, and negative values argue in favor of the presence of
oxidized surface waters during deposition (e.g., Towe, 1991; Bau
and Dulski, 1996; Shields and Stille, 2001; Pi et al., 2013). The Ce/
Ce* values are <1.0 in all Meishan samples, with a slight increase
in the LPME horizon. In Guryul Ravine, Ce/Ce* values are <1.0 in
pre-LPME samples and show a discrete negative shift at the
LPME horizon. However, it undergoes an increase across the PTB
horizon and a more pronounced one at the ETME horizon. These
patterns suggest that anoxic conditions prevailed across the
LPME and dysoxic to oxic conditions across the PTB and
ETME in Meishan. In Guryul Ravine, dysoxic to oxic
conditions dominated across the LPME horizon and changed
to anoxic around Bed 49 and across the ETME horizon. The Eu/
Eu* values <1.0 in all Meishan samples show a marked negative
peak at the LPME. The Pr/Pr* values are >1.0 with a positive peak
at the LPME horizon. In the Guryul Ravine section, Eu/Eu* > 1.0
at the LPME horizon and Pr/Pr* > 1.0 in beds 43 through 46.
Positive Eu/Eu* values dominate above the PTB horizon and
probably reflect clay mineral provenance rather than a
hydrogenous signal, according to Bhandari et al. (1992).
To better visualize the Ce behavior by eliminating the
influence of La, the Ce/Ce* and Pr/Pr* ratios were plotted
against each other (Figure 3C), following Bau and Dulski
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(1996). The Meishan section displays a slightly different
pattern as compared to the Guryul Ravine section
(Figure 3C). Several pre-LPME samples and those from the
extinction interval and the PTB and ETME horizons plot
within the “true negative Ce anomaly” box. Instead, all
samples from the Guryul Ravine section plot between the
negative and positive anomaly boxes, except a sample from
the ETME horizon, which plots within the “true positive
anomaly” box. In a Ce anomaly [log(Ce/Ce*)] vs. Nd (ppm)
plot (Figure 3D), following (Wang et al., 2014; modified from
Elderfield and Pagett, 1986; Wright et al., 1987), all samples from
the Guryul Ravine section and most samples from Meishan plot
within the anoxic field. Five samples from Meishan plot in the
oxic field very close to the oxic-anoxic field boundary (0.0
and –0.1).
The highest Nd values from the Meishan section are >60 ppm
and all in the anoxic field (bed 26) in Figure 3D, while the lowest
Ce anomaly value is observed in the LPME horizon, which falls
within the oxic field. This argues in favor of an abrupt change
from oxic to anoxic redox conditions across the LPME horizon in
Meishan. At Guryul Ravine, the two highest Nd values (46 and
39 ppm) are observed in samples from the top of bed 46 and top
of bed 48. The occurrence of framboidal pyrite in bed 49 supports
the transition from oxic to anoxic conditions in beds 48–49.
Carbon Isotope Chemostratigraphy
Meishan Section
The δ13Ccarb and δ13Corg patterns for Meishan section are shown
in Figure 4. Forty samples (Table 2) from a previous study (Jin
et al., 2000; Sial et al., 2020a) have also been used to complete the
δ 13Ccarb profile. Further data from the extinction interval (Chen
et al., 2005) have been added to enhance the resolution of the
δ 13Ccarb curve across the PTB.
The δ13Ccarb values in the 24–29 bed interval varied from −1.2
to +4.7‰, prevailing positive values. All samples from the
Changhsing Formation yielded positive δ 13Ccarb values, except
the topmost beds 24 and 25 (LPME) with a value of −1.1‰. A
positive shift >1‰ in bed 27a (Clarkina taylorae Zone) coincides
with a change from anoxic to dysoxic/oxic conditions. At the PTB
horizon (bed 27c) a δ 13Ccarb value of +1.1‰ was measured, and
+0.9‰ at the base of bed 28 (ETME horizon). In bed 29, all
δ 13Ccarb values are positive and gradually increase upwards (beds
30–40 are not shown).
The δ13Corg curve in Figure 4 is from Sial et al. (2020a) with
additional 12 samples analyzed in this study, showing δ 13Corg
varying between −30.5 and −25.7‰. Covariation of δ13Ccarb and
δ 13Corg is observed at Meishan (Figure 4). A discrete positive shift
of δ 13Ccarb in bed 27 matches a more marked positive shift of the
δ 13Corg values, coinciding with a change from anoxic to dysoxic/
oxic conditions.
A careful examination of the geochemistry, mineralogy, and
cathodoluminescence of beds 24–62, spanning the PTB in the
Meishan D section, led Li and Jones (2017) to conclude that the
δ 13C signals were a mixture of primary and diagenetic signals.
They cautioned that in dolomite-bearing samples, diagenesis
might have contributed ∼1.7‰ to the negative δ13C shift.
Besides this diagenetic overprint, Kaiho et al. (2006) explained
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a −1‰ shift of δ 13C in the Meishan section, within ∼13 ka before
the LPME, by possible thermal dissociation of methane hydrates.
Subsequent mixing of an anoxic stratified ocean led to another
–1‰ δ 13C shift within ∼5 kyr of the extinction. According to
these authors, low ∆13C (∆13C δ 13Ccarb − δ 13Corg) values in beds
27–29 indicate that phototrophic sulfur bacteria may have
flourished and lasted over 150 ka after the LPME in the lowlatitude Paleotethys.
Coeval sections may allow testing the authenticity of major
δ 13C shifts, yet diagenetic signals may be globally synchronous
(Fantle et al., 2020). In any case, absolute values of major δ 13C
changes at Meishan, Guryul Ravine, or any other section need to
be considered cautiously.
In summary, the C-isotope curves from Meishan and Guryul
Ravine highlight the following features: a) a pronounced negative
δ 13C excursion at the LPME horizon in both sections b) a negative
δ 13C excursion at the PTB horizon, c) a discrete positive δ 13Ccarb
and δ13Corg shift (of unclear origin) in the LPME–PTB interval, d)
at Guryul Ravine the δ13Ccarb return to positive values above the
PTB, while at Meishan they tend to stay steady, e) a modest
response of δ13C values to environmental effects associated with
the ETME as compared to those observed at the LPME, and f) a
sharp increase in δ13Ccarb values, a few meters above the ETME
(e.g., Song et al., 2013).
Mercury Chemostratigraphy
Mercury (Hg) enrichments across some chronological
boundaries can be the result of volcanogenic Hg loading onto
the environment (e.g., Sanei et al., 2012; Sial et al., 2013; Sial et al.,
2016; Sial et al., 2019; Sial et al., 2020a; Sial et al., 2020b; Grasby
et al., 2015a; Grasby et al., 2017; Grasby et al., 2019; Grasby et al.,
2020; Percival et al., 2015; Percival et al., 2017; Percival et al.,
2018; Font et al., 2016; Thibodeau et al., 2016; Thibodeau and
Bergquist, 2017; Charbonnier et al., 2017; Meyer et al., 2019; Shen
et al., 2019a,b,c; Them et al., 2019; Kwon et al., 2019, and
references therein). This framework allows examining the
relationship between LIP activity and periods of extreme
environmental turnover. The application of Hg for this
purpose requires its normalization to an element representing
the dominant host phase of Hg―generally total organic carbon
(TOC) for the organic fraction, but occasionally total sulfur (TS)
or aluminum (Al) for the sulfide or clay fractions, respectively
(e.g., Shen et al., 2020).
The Hg/TOC stratigraphic curve (Figure 4) in Meishan shows
three peaks, a prominent one around the LPME (peak I, top of
Bed 24), a second one (peak II, Bed 25), and a third one (peak III)
at the top of Bed 26. The Hg/TS curve, on the contrary, is very
monotonous and characterized by only a single peak in Bed 26
(Hindeodus changxingensis biozone), in an interval with
framboidal pyrite (i.e., anoxic/euxinic conditions). The Hg/Al
curve shows higher values in pre-LPME samples (Bed 24), lower
values in the extinction interval, and a prominent peak at the
LPME horizon, coinciding with the Hg/TOC peak I and the
negative δ 13Ccarb shift. The Hg/TOC, Hg/S, and Hg/Al curves and
correlation coefficients between Hg and TOC (∼0.50), S (0.11),
and Al (0.25) suggest that organic matter was probably the main
Hg-host phase. Only in the anoxic interval of bed 26 the role of
framboidal pyrite is evidenced as an additional Hg-host phase.
Aluminum concentrations are higher in the extinction interval
samples (2.4–13.57%, in beds 25–29) than those from the preLPME (0.11–0.4%, bed 24). Although total Hg is higher in
samples from the extinction interval, Hg/Al behaves
oppositely, with higher values in pre-LPME samples, except
for the Bed 24 sample with anomalously high Hg, which
causes a prominent Hg/Al peak.
The Hg/TOC curve for Guryul Ravine section (Figure 5)
shows lower values in pre-LPME samples (beds 43–46, Zewan
Formation), the peak I at the LPME horizon, peak II in bed 49
(Khunamuh Formation), peak III at the PTB horizon (base of Bed
52) and a larger peak IV at Bed 53 (ETME horizon). The Hg/S
Guryul Ravine Section
The first δ 13Ccarb curves for the Guryul Ravine section were
reported by Baud and Magaritz (1988) and Baud et al. (1989,
1996).
Baud et al. (1989, 1996) reported the first δ13Ccarb curves for
the Guryul Ravine section. Korte et al. (2010) provided a new
δ 13Ccarb chemostratigraphy for this section. These authors
examined the role of massive volcanism on the isotopic
composition of the ocean and atmosphere at the PTB. The
high-resolution δ13Ccarb curve in Figure 5 is from Sial et al.
(2020b). The high-resolution δ 13Ccarb variation curve starts from
bed 25 (the topmost bed of the Member B, Zewan Formation) and
ends in bed 53 (Member E3, Khunamuh Formation) and has an
overall similarity with the δ 13Ccarb curve of Korte et al. (2010).
However, the extinction interval being our main target, only
δ 13Ccarb values from beds 43 through 46 (part of Member D of the
Zewan Formation) and beds 47 through 53 (Members E1 and
part of Member E2 of the Khunamuh Formation) are listed in
Table 1 and displayed in Figure 5.
Whereas Member D yielded ∼0.0‰ to moderately positive
δ 13Ccarb values (+0.1–+3.1‰), the LPME and PTB horizons are
characterized by strong negative excursions (−4‰). A positive
δ 13Ccarb shift is observed at the transition from bed 50 to 51 and
another one at the base of bed 53 (praeparvus and isarcica zones)
that, according to Huang Q. et al. (2019), was deposited during a
stage of widespread oceanic oxygen deficiency. This δ13Ccarb shift
is, perhaps, the record of the ETME event at Guryul Ravine. In
bed 53 and upsection, positive δ 13Ccarb values dominate and
progressively rise to +3‰ (Member E3, not shown in Figure 5).
The δ 13Corg curve (Figure 5) shows values between –28 and
–24‰ for samples from the Zewan Formation, consistent with
organic matter of both marine and terrestrial sources. Algeo et al.
(2007) argued in favor of a dominantly marine origin based on
low molar C/N ratios (5:1–10:1) and δ 15N values reported by
White (2020). The first negative excursion in both the δ13Ccarb
and δ13Corg curves is observed in bed D45, ∼6 m below the PTB
and ∼3.3 m below the contact between the Zewan and Khunamuh
formations. A positive δ 13Corg shift at the LPME horizon
coincides with a positive excursion in the δ 13Ccarb values (base
of bed 47). A discrete positive δ 13Corg shift in bed 49 (Member E1)
is within the LPME–PTB interval (C. meishanensis–H.
praeparvus Zone). Framboidal pyrite in this horizon was first
mentioned by Brookfield et al. (2013).
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Mercury Isotopes
curve shows a monotonous behavior, except for two peaks at the
base of Bed 47 (at the contact of Zewan and Khunamuh
formations), and in the framboidal pyrite-bearing Bed 49. The
Hg/Al curve is also monotonous as a whole, with three discrete
peaks, at the LPME, PTB and ETME horizons.
The behavior of Hg/TOC, Hg/TS and Hg/Al stratigraphic
curves and the calculated correlation coefficients of Hg vs. TOC
(0.27), Al (0.20), and TS (∼0.01) suggest that TOC has been the
main Hg-host phase. Aluminum concentrations increase slightly
beween the top of Bed 46 and Bed 51 where shales were deposited
at the PTB in Guryul Ravine during a transgressive event. A
discrete Hg/Al peak records the change to more Al-enriched
cabonate rocks or shales, but on the whole, clays played a less
significant role as the Hg-host phase, as compared to the organic
matter.
It is clear from Figure 4 and the above discussion that Hg/
TOC ratios show more intense environmental changes across the
PTB than the carbon isotope (δ13Ccarb, δ 13Corg)
chemostratigraphy. These differences may be related to the
geochemistry of Mercury and carbon. In short, Mercury is a
chalcophile (sulfur-loving) element and biologically highly active.
Due to its relatively high vapor pressure, it may be mobilized
tectonically and in the gas phase (e.g., Fitzgerald and Lamborg,
2005). In the surface environment, Hg is a very active element
(i.e., biologically active, redox-active, volatile) and has a strong
affinity to organic matter. Mercury is therefore more easily
impacted by source input changes (volcanic, tectonic) and
environmental and climate changes as compared to carbon. As
such, the Hg/TOC ratio can be regarded as a sensitive and
promising proxy for environmental and ecological changes.
Carbon isotope chemostratigraphy may show indistinct
patterns as compared to Mercury. Both δ 13Ccarb and δ 13Corg
values are affected by several primary factors (e.g., atmospheric
CO2 concentration and δ 13C value, pH and salinity of the
seawater, biomass productivity, organic matter type) as well as
post-depositional processes (e.g., recrystallization and
neoformation of carbonate phases during early and burial
diagenesis, rock interaction with hydrothermal and
metamorphic fluids, preferential preservation of organic matter
during diagenesis, organic matter thermal maturity).
To underline the application of Hg/TOC profile as a tool for
stratigraphic correlation, Figure 6 depicts a Hg/TOC NW-SE
transect diagonally across the Pangea supercontinent at 252 Ma,
passing through seven well-known PTB sections (from Ursula
Creek, Canada, to Hovea-3, Australia) plus Meishan and Guryul
Ravine (based on Sial et al., 2020a). The LPME and ETME
horizons were kept as they appear in the original figures, and
the PTB was used as a datum. The largest Hg/TOC spike is
observed at the LPME horizon in most of these sections, a smaller
one at the PTB and a more discrete one at the ETME horizon.
This transect reveals that, except for the Idrijca section, all the
other sections exhibit a discrete Hg/TOC peak of unknown origin
between the LPME and the PTB. This peak may record a global
event as it is also observed outside of the Tethyan realm. More
importantly, Figure 6 clearly shows how powerful the Hg/TOC
ratios are as a stratigraphic correlation tool, mainly when a timedefined level had Hg input from LIP activity.
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The Mercury source can be potentially traced with the help of Hg
isotopes as they can undergo large mass-dependent (MDF) and
mass-independent (MIF) fractionations in nature (Blum et al.,
2014). These isotopes can also help interpret Hg and Hg/TOC
trends (e.g., Thibodeau et al., 2016; Gong et al., 2017; Grasby et al.,
2017; Wang et al., 2018). It is necessary to consider that the Hg
isotopes MDF (hereafter Hg-MDF) can be triggered or modified
by physical, chemical, or biological reactions. These parameters
need to be considered when tracing Hg in the environment. As
Thibodeau et al. (2016) pointed out, the Hg isotopes MIF
(hereafter Hg-MIF) are more useful for source identification as
they are not affected by post-depositional processes. The nearzero Hg-MDF values indicate direct volcanic input, while more
negative values indicate terrestrial sources related to organic
matter heating or enhanced erosion (Yin et al., 2016).
According to Wang et al. (2018), Hg-MIF (Δ199Hg) with
sustained positive values indicates a predominant atmosphericderived signature of volcanic Hg in deep-shelf settings, while
nearshore environments display a negative Δ199Hg signature,
interpreted to be linked to terrestrial Hg sources.
Twenty-four samples were analyzed for the Hg isotopes, nine
from the Meishan D section and 15 from the Guryul Ravine
section (Table 2). The samples are predominantly from the
extinction interval (LPME–ETME) and yielded high Hg/TOC
values. Besides, 20 Hg-isotope analyses from the Meishan D
section reported by Grasby et al. (2017) were plotted together
with our results in Figures 7A, 8A. All the Hg-MDF values,
reported here as δ 202Hg, are negative in both sections. In Meishan
D, almost all the δ 202Hg values are in the −2.00–−0.50‰ range,
amongst which two samples are between −1.00 and −0.50‰
(Table 2) and only two values are lower than −2.00‰. Almost all
pre-LPME samples of this section display positive Δ201Hg (MIF)
values and tend to be negative at the extinction interval. In the
Guryul Ravine section, 60% of the samples have δ 202Hg values
between −2.00 and −1.50‰, and six of them are below −2.00‰.
The pre-LPME samples generally have positive Δ201Hg values,
while negative values prevail in the extinction interval.
In the Meishan D section, over 85% of the samples plot in the
“volcanic-emission box” in the δ202Hg vs. Δ201Hg diagram
(Figure 7A; Sial et al., 2016). Two samples fall in the
“volcanic emission and chondrite box” and one (top of Bed
24E) in the “sediment, soil and peat” box. Three horizontal
trends are shown in Figure 7A: trend I formed by Pre-LPME
samples, trends II and III by samples from the extinction interval.
With positive Δ201Hg values, trend I is delineated by samples
from Bed 24 (data from Grasby et al., 2017, and from the present
study) in which one sample from the top of Bed 24E falls within
the “sediment, soil and peat” box. Trend II (samples from Bed 25)
and trend III (Bed 26) show negative Δ201Hg values. Only one
sample from the LPME horizon plots in trend I and a sample
from the ETME horizon in trend II.
Most probably, Hg in the samples from Meishan right below
the LPME (Bed 24) and within the LPME–PTB (beds 25 and 26)
originated from the same volcanic source, with negligible or no
Δ201Hg (MIF) fractionation during atmospheric transport. The
range of measured Δ199Hg values (−0.10–0.00) within this
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In trend II, Δ201Hg varies from zero to slightly negative values and
concentrates in the “volcanic emission box.” It is likely that these
two trends have been influenced by a major eustatic event
initiated at the base of Bed 46. This event, according to
Brookfield et al. (2003), was characterized by an initial
regression (Bed 46) before a transgression (beds 47–50). Algeo
et al. (2007) have favored transgression alone, with sea-level rise
initiated at the base of Bed 46 and reaching maximum flooding
conditions around the PTB horizon. The sea-level changes
probably influenced the shape of the Hg-isotope trends in
Figure 7B: trend I (Bed 46) during the regression suggested
by Brookfield et al. (2003), and trend II (beds 47–52, extinction
interval) by the subsequent sea-level rise during the transgression.
The Δ199Hg vs. total Hg plot shows two trends for the Guryul
Ravine section (Figure 8B). Trend I, defined by pre-LPME
samples (bed 46) with positive Δ199Hg values and an almost
horizontal trend II defined by samples from beds 47–53
(extinction interval) with slightly positive Δ199Hg values, except for
two samples. It is possible that in trend I, the Δ199Hg values in Bed 46
samples represent mixing of terrigenous and volcanic sources of Hg,
during a major eustatic sea-level variation. It is important to note that in
a study of Hg anomalies from shallow and deep-water depositional
environments across the LPME in South China, Wang et al. (2018)
observed that: a) positive Δ199Hg values predominate in deep shelf
settings, as atmosphere-derived signature of volcanic Hg (e.g., Daxiakou
and Shangsi sections), and b) negative Δ199Hg values predominate in
nearshore environments at Meishan D and are likely associated with the
terrestrial Hg sources. Samples in trend I (Figure 8A) probably mark the
time interval when atmospheric-derived Hg was deposited in a
deepening-upward sequence from shallow (negative Δ199Hg) into a
deep environment (positive Δ199Hg). This stage was followed by a
gradual deepening of the basin as transgression progressed, and Δ199Hg
values in trend II are all positive. Besides, the PTB horizon displays a
positive value, higher than values in trend II, as it represents the
maximum flooding in this transgression.
interval is within experimental error of zero, supporting a
significant influx of volcanic Hg in this time interval.
Moreover, small positive Δ201Hg values observed in bed 24
favors long-term atmospheric transport and Hg loading to the
environment by the STLIP magmatism. MDF of Hg isotopes
seem to have been affected by local fractionation in trend I,
pushing δ 202Hg values towards more negative values (e.g., sample
from the top of the Bed 24E) and, to a lesser extent, in trend II
(sample MS 25de from Bed 25; Table 2) due to terrestrial Hg
influx. Alternatively, a coal seam intruded by a contemporary sill
complex during Stage 2 of the STLIP magmatic activities (Burgess
et al., 2017) could also explain this feature. It is unlikely that postdepositional processes had any effect on the Hg-MIF signatures.
Therefore, we can safely assume that the Δ201Hg values in the
three observed trends in Figure 7A represent primary signals,
originated by the Δ201Hg STLIP signatures.
The Δ199Hg curve for Meishan is characterized by positive
values in pre-LPME samples (bed 24), a marked negative shift at
the LPME horizon, and predominantly negative values over the
extinction interval (beds 26 and 27) returning to positive values at
the top of bed 29 (Figure 4). A similar pattern has been reported
by Wang et al. (2018) using samples from Grasby et al. (2017),
shown in Figure 4 as gray triangles along the Δ199Hg curve. The
Δ199Hg (odd-MIF) curve of Guryul Ravine section displays
positive values in pre-LPME samples (beds 45 and 46) and at
the LPME horizon. A negative shift is observed in the extinction
interval (beds 48 and 51) with a positive shift in beds 49 and 50. A
discrete peak is present at the PTB horizon, and positive values
predominate upsection.
In the Δ199Hg (MIF) vs. total Hg plot of the Meishan D section
five trends are observed (Figure 8A). The trend I is defined by
pre-LPME samples (Bed 24), while samples from the extinction
interval (beds 25 and 26) represent the curved trends II and III.
The trend I is consistent with Hg deposition before the sea-level
fall recognized in several studies (Wu et al., 2014 and references
therein) at about 0.8 m below the PTB at Xiushui PTB section,
Jiangxi Province, South China. This level corresponds to the
beginning of the Hindeodus changxingensis zone, correlatable to
Hindeodus typicalis zone in the Meishan section. It is possible that
the trend I reflects a deep-shelf setting deposition of atmospheric
volcanic Hg deposition just before the above-mentioned sea-level
drop. Trend II (beds 25 and 26) is consistent with a twocomponent mixing between terrigenous and volcanic-source
Hg deposition in progressively shallower shelf settings during
sea-level drop. This trend is formed by samples with both positive
and negative Δ199Hg (MIF) values, while trend II (Bed 26)
includes only samples with negative Δ199Hg (MIF) values. Up
section, samples from bed 27 with negative Δ199Hg (MIF) values
depict trend IV, and three aligned samples from bed 29 (positive,
zero, and negative Δ199Hg values) suggest an additional trend
(trend V).
The δ 202Hg vs. Δ201Hg plot shows two trends for the Guryul
Ravine section (Figure 7B). Pre-LPME samples from Bed 46
define the trend I, and samples from the extinction interval in
beds 46 through 53 form trend II. In trend I, all samples display
positive Δ201Hg values, and two of them fall in the “volcanic
emission box” and three within the “sediment, soil and peat” box.
Frontiers in Earth Science | www.frontiersin.org
CONCLUSION
The main conclusions of this study are listed below:
1. Selected trace elements, elemental ratios, and REE + Y confirm
varying paleoredox conditions in the Meishan D and Guryul
Ravine sections.
2. Sea-level fluctuations across the PTB transition in these two
sections controlled the detrital supply, as shown by several
detrital input proxies.
3. Both sections recorded more intense environmental changes
at the LPME and ETME horizons, reflected in C-isotope and
Hg chemostratigraphic curves.
4. The Hg/TOC spikes coincide with a strong negative δ 13C shift
at the LPME horizon in both sections, probably synchronously
with the initiation of the STLIP intrusive stage of magmatism
(stage 2).
5. The main Hg/TOC spike in Meishan occurs at the LPME
horizon; in Guryul Ravine it is at the ETME horizon, with less
pronounced spikes at the PTB and LPME horizons.
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June 2021 | Volume 9 | Article 651224
Sial et al.
Hg Chemostratigraphy Permian-Triassic Boundary
AUTHOR CONTRIBUTIONS
6. A discrete Hg/TOC peak in Meishan (peak III) marks the change
from anoxic to oxic paleoredox conditions (base of bed 27). An
identical discrete spike (peak II) in Guryul Ravine at the base of bed
49 marks a change from dysoxic/oxic to anoxic conditions. The
origin of this spike within the LPME–PTB interval, also observed
in other PTB sections, remains an open question.
7. Normalization of Hg concentration to TOC, total sulfur or
total aluminum in the studied samples allowed to infer that
organic matter is the main Hg-host phase, compared to Hg
hosted in sulfides or adsorbed on clay minerals.
8. Hg chemostratigraphy reveals more intense environmental
changes at PTB than carbon isotope chemostratigraphy.
9. In particular, the Hg/TOC values represent a robust tool in
stratigraphic correlation, as shown here for PTB sections in a
transect across the Pangea Supercontinent.
AS: writing and interpretations. JC: analyses of Hg and Hg
isotopes. CK: sampling collection, trace elements, Total
Organic Carbon (TOC), and Rare Earth Elements (REE)
analyses. LL: Hg geochemistry interpretation. MP: field work
and sample collecting. JS: trace element chemostratigraphy.
JS-T: samples from PTB sections. CG: text writing. VF:
sample preparation and text writing. JB: figures and
interpretation. NP: figures and interpretation. PR: figures
and interpretation.
FUNDING
The study was partially supported by grants to AS (CNPq 407171/
2018-5; FACEPE APQ-1073-1.07/15), to LL (CNPq INCTTMCOcean 573.601/2008-9, CNPq576.601/2009-1), to VF
(FACEPE APQ1738-1.07/12) and NSFC grants to JC (No.
41625012, U1612442, 4196114028, 41830647). MP thanks
Ahsan Wani for the support and assistance during
fieldwork and sampling at the Guryul Ravine and CK
thanks the Danish Natural Science Research Council for
providing financial support for sampling in China and to
QuanFeng Zheng for leading him in the field in Meishan
(Grant Number 11-103378).
Finally, Hg isotopes serve to refine the potential Hg source and
associated paeoenvironmental processes. The pre-LPME samples
from Meishan and Guryul Ravine exhibit positive Δ199Hg (MIF)
values. In the LPME–PTB interval, negative values dominate and
positive values are seen again above the ETME. These Δ199Hg patterns
and correlation with total Hg reflect the influence of water depth in the
deposition of atmospheric volcanic Hg at both sections, in turn,
dependent upon sea-level variations (regression-transgression).
Alternatively, these patterns may represent the three main stages of
the dominant style of STLIP magmatism relative to the timing of mass
extinction across the PTB. In stage 2 (LPME–PTB interval), a complex
of sills intruded coal-bearing strata. Pre-LPME samples have positive
Δ201Hg (MIF) values, but negative values dominate in the extinction
interval. Most samples from Meishan and Guryul Ravine fall within
the “volcanic-emission box” in the Δ202Hg (MDF) vs. Δ201Hg (MIF)
plot. The Δ 201Hg (MIF) values suggest primary signals controlled by
STLIP Δ201Hg signatures.
ACKNOWLEDGMENTS
We are grateful to Olga Nielsen of De Nationale Geologiske
Undersøgelser for Danmark og Grønland—GEUS. We
acknowledge also J. Bojesen-Koefoed and GEUS for TOC, TC
and Rock-Eval pyrolysis measurements, as well as O. Nielsen
(GEUS) and B. Petersen (IGN) for running ICP-MS and carbon
isotope analyses, respectively. We thank Clemens Ullmann
(University of Exeter, United Kingdom) for evaluating the
statistics of the ICP-MS analyses. Comments from two
anonymous reviewers have helped in improving the manuscript.
This is the NEG–LABISE scientific contribution n. 300.
DATA AVAILABILITY STATEMENT
The original contributions presented in the study are included in
the article/Supplementary Material, further inquiries can be
directed to the corresponding authors.
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