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Article https://doi.org/10.1038/s41467-023-39427-z
Recurrent photic zone euxinia limited ocean
oxygenation and animal evolution during
the Ediacaran
Wang Zheng
1
, Anwen Zhou
1,2
,SwapanK.Sahoo
3
,MorrisonR.Nolan
4
,
Chadlin M. Ostrander
5,6
,RuoyuSun
1
, Ariel D. Anbar
7,8
, Shuhai Xiao
4
&
Jiubin Chen
1
The Ediacaran Period (~635–539 Ma) is marked by the emergence and diversi-
fication of complex metazoans linked to ocean redox changes, but the pro-
cesses and mechanism of the redox evolution in the Ediacaran ocean are
intensely debated. Here we use mercury isotope compositions from multiple
black shale sections of the Doushantuo Formation in South China to reconstruct
Ediacaran oceanic redox conditions. Mercury isotopes show compelling evi-
dence for recurrent and spatially dynamic photic zone euxinia (PZE) on the
continental margin of South China during time intervals coincident with pre-
viously identified ocean oxygenation events. We suggest that PZE was driven by
increased availability of sulfate and nutrients from a transiently oxygenated
ocean, but PZE may have also initiated negative feedbacks that inhibited oxygen
production by promoting anoxygenic photosynthesis and limiting the habitable
space for eukaryotes, hence abating the long-term rise of oxygen and restricting
the Ediacaran expansion of macroscopic oxygen-demanding animals.
The Ediacaran Period (635–539 million years ago, or Ma) was a pivotal
timeframe in the evolution of life on Earth, evidenced by the initial
appearance of macroscopic multicellular eukaryotes in the fossil
record (i.e., the “Ediacara biota”) and its subsequent decline1. The rise
and decline of the Ediacara biota are among the biggest mysteries in
the evolution of life. Numerous studies have linked the rise of mul-
ticellular eukaryotes and the diversification of the Ediacara biota to
ocean oxygenation2,3. Likewise, the decline of the Ediacara biota has
been linked to a drop in ocean oxygen levels during the terminal
Ediacaran Period4,5. However, the causal relationships between bio-
logical innovation and ocean redox during this period remain deba-
ted because of disparate ocean redox conditions reconstructed by
different paleoredox proxies6. Redox-sensitive element (RSE)
enrichments and carbon and sulfur isotope variations found in the
sedimentary record are interpreted as evidence of widespread oxic
conditions, and perhaps even ventilation of the deep ocean, at least
episodically during the Ediacaran2,3,7,8. Yet, Fe speciation data indi-
cate persistently anoxic deep waters throughout most of the
Ediacaran9,10. In recent years, a growing body of evidence supports a
redox-stratified Ediacaran ocean, with some continental shelf mar-
gins overlain by a dynamic mid-depth euxinic water mass (“euxinic
wedge”)11,12. In this scenario, early animals could have thrived in oxic
shelf settings even as deep oceans remained anoxic13,14. It was even
suggested that surface ocean environments that met the oxygen
requirements of the earliest metazoans (0.5–4.0% of present atmo-
spheric levels) were already established well before the Ediacaran15,
leading to the proposal that the radiation of animals was not pro-
hibited by oxygen availability. Thus, additional information about
Received: 31 January 2023
Accepted: 12 June 2023
Check for updates
1
School of Earth System Science, Institute of Surface-Earth System Science, Tianjin University, Tianjin 300072, China.
2
Department of Earth, Ocean and
Atmospheric Science and National High Magnetic Field Laboratory, Florida State University, Tallahassee, FL 32306, USA.
3
Equinor US, Houston, TX, USA.
4
Department of Geosciences, Virginia Tech, Blacksburg, VA 24061, USA.
5
Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic
Institution, Woods Hole, MA 02543, USA.
6
Department of Geology and Geophysics, University of Utah, Salt Lake City, UT 84112, USA.
7
School of Earth and
Space Exploration, Arizona State University, Tempe, AZ 85287, USA.
8
School of Molecular Sciences, Arizona State University, Tempe, AZ 85287, USA.
e-mail: swas@equinor.com;jbchen@tju.edu.cn
Nature Communications | (2023) 14:3920 1
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ocean redox conditions and their changes could help to resolve these
complexities.
In this study, we refine the redox evolution in the Ediacaran ocean
using mercury (Hg) stableisotopes. In particular, we focus on evidence
for the presence of anoxic and H
2
S-rich waters in the marine photic
zone, a condition referred to as photic zone euxinia (PZE)16,17. Different
from the previously proposed “euxinic wedge”that refers to a mid-
depth euxinic water mass11,12, PZE emphasizes the shallowing of the
euxinic water mass into the photic zone, which is particularly detri-
mental to shallow marine inhabitants and has been proposed as a
potent kill mechanism during almost all Phanerozoic mass extinction
events17 as well as a key factor responsible for the evolutionary stasis
during the mid-Proterozoic18. The shelf areas are believed to have been
major habitats for early animals19, and thus the development of PZE
would have shrunk the oxygenated habitats and possibly contributed
to the delayed rise and eventual decline of the Ediacara biota. More-
over, PZE is typically found to be driven by enhanced nutrient influx to
the ocean20, which is exactly the mechanism proposed to trigger the
oxygenation events in the Ediacaran ocean7,8. Therefore, there could
be a potential link between ocean oxygenation and PZE, and thus the
investigation on PZE may provide insights into the cause and con-
sequence of ocean oxygenation in the Ediacaran Period.
Mercury (Hg) isotopes in sedimentary rocks are an emerging
proxy for tracing paleoenvironmental changes such as large volcanic
emission21,aswellasPZE
16. A previous study has shown that marine
sediments deposited under PZE display distinct negative excursions of
mass independent isotope fractionation (MIF) of Hg relative to non-
PZE conditions,demonstrating that Hg isotopes are a promising proxy
for changes in ocean redox conditions16. Here we report Hg con-
centration and isotope ratios of the Ediacaran Doushantuo Formation
in South China (Fig. 1). We studied four black-shale dominated sec-
tions, including Weng’an (WA, shallow shelf), Taoying (TY, upper
slope), Wuhe (WH, lower slope), and Yuanjia (YJ, ba sin) sections, which
represent a shelf-to-basin transect. The Doushantuo Formation
(~635–560 Ma) is known for its exceptionally well-preserved fossil and
geochemical record, and has been extensively studied to reconstruct
the Ediacaran ocean environments22. Previous studies on the same
suite of black shale sections reported evidence for three ocean oxy-
genation events (OOEs) at ~635 Ma, ~580 Ma, ~565 Ma based on a
variety of redox proxies (i.e., Fe speciation, RSE and δ34S
pyrite
)7,8. Var-
ious metal isotopes (δ98Mo, δ53Cr, and ε205Tl) have also been measured
in these samples23–25. In the context of the published data, our Hg
isotope data provide strong evidence for recurrent PZE as well as its
impact on the oceanic redox evolution and biological innovation
during this critical period.
Results and discussion
Hg concentration and isotope compositions
In WH, thetotal Hg concentrations (THg) are >100 ppb except for only
two samples. Member IV, which correlates with the previously pro-
posed OOE interval, shows notably higher THg than Member II and III
(Fig. 2). To constrain the mechanism of Hg enrichment in sediments,
THg is often normalized tomajor host phases of Hg, including organic
matter (proxied by total organic carbon, TOC), sulfide (proxied by
pyrite S, S
py
) and clay minerals (proxied by Al concentration)26;the
ratios of Hg/TOC, Hg/S
py
and Hg/Al are all significantly elevated in
the Member IV OOE interval of WH (Fig. 2). The TY section, where only
the three OOE intervals were analyzed, shows in general the highest
THg among all sections and an increasing trend of THg and Hg/TOC
from Member II to IV. YJ and WA sections were only analyzed for the
MemberII OOE interval.In YJ, THg and normalized ratios (Hg/TOC, Hg/
S
py
and Hg/Al) show a transient elevation at the lowermost of the
section, and decreases up-section. In WA, the variations of THg and
normalized ratios show no apparent trend. The THg of different
intervals is correlated with different host phases, and these correla-
tions are described in details in Supplementary Text S1 and Supple-
mentary Fig. S1.
The Δ199Hg values (representing MIF of odd isotopes) of WH show
cyclic variability, with distinct negative excursions at the lowermost
Member II and III, as well as Member IV (down to −0.10‰in the low-
ermost Member II), coinciding with three previously proposed OOE
intervals8,andgradualpositiveshifts(upto0.19‰) between OOEs
(Fig. 2). To facilitate discussion, we label the three intervals that show
negative excursions of Δ199Hg in Member II, III and IV of WH and the
b
Tidal Flat
Slope
Chaoyang
Jiangchuan
3
2
1
1
4
Zunyi
Guiyang
Chongqing
Yichang
12
3
4
North China
Platform
Shanghai
Chengdu Yichang
Chongqing
Changsha Nanchang
Guiyang
Kunming
Nanning Guangzhou
Hongkong
Vietnam
Taibei
Wuhan
a
11
2
3
4
Weng’an
Taoying
Wuhe
Yuanjia
Lantian
Chengjiang
Fig. 1 | Paleogeographic location of the study sections in the Yangtze Platform
duringthe depositionof the DoushantuoFormation in South China.The region
marked by the red rectangle in panel acorresponds to panel b. The numbers in
yellow circles are the study sections: Weng’an (WA), Taoying (TY), Wuhe (WH),
Yuanjia (YJ). Reprinted from Jiang, G., Shi, X., Zhang, S., Wang, Y. & Xiao, S. Strati-
graphy and paleogeography of the Ediacaran Doushantuo Formation (ca.
635–551 Ma) in South China. Gondwana Res. 19, 831–849 (2011), with permission
from Elsevier.
Article https://doi.org/10.1038/s41467-023-39427-z
Nature Communications | (2023) 14:3920 2
Content courtesy of Springer Nature, terms of use apply. Rights reserved
I
II
III
IV
0 1000 2000
0 300 600
0.0 0.4 0.8
0.0 0.4 0.8
0 500 1000
0 500 1000
0 100
02040-3 -2 -1 0
-0.1 0.0 0.1 0.2 -40 -20 0 20 0 4000 8000
-0.1 0.0 0.1
Hg (ppb)
Hg/TOC
Fe
HR
/Fe
T
Fe
PY
/Fe
HR
Hg/S
py
Hg/Al
Mo (ppm)
U(ppm) G
202
Hg (‰)
'
201
Hg (‰)
'
199
Hg (‰)
G
34
S
pyrite
(‰) P(ppm)
0 1500 3000
P/Al
'
200
Hg (‰)
a. WH
E3
E2
E1
I
II
III
IV
0 1500 3000
0 800 1600
0.0 0.4 0.8
0.0 0.4 0.8 1.2
0 4000
0 150 300
0 100 200
10 20
-0.5 0.0 0.5 -0.1 0.0 -0.1 0.0 0.1
Hg (ppb)
Hg/TOC
Fe
HR
/Fe
T
Fe
PY
/Fe
HR
Hg/S
py
Hg/Al
Mo (ppm)
U(ppm)
G
202
Hg (‰) '
201
Hg (‰)
'
199
Hg (‰)
'
200
Hg (‰)
b. YJ
E1
I
II
III
IV
0 2000 4000
0 800 1600
0.0 0.4 0.8
0.0 0.4 0.8
0 50000
0 1000 2000
0 100 200
040-2 -1 0
-0.1 0.0 0.1 0.2 -0.1 0.0 0.1
Hg (ppb)
Hg/TOC
Fe
HR
/Fe
T
Fe
PY
/Fe
HR
Hg/S
py
Hg/Al
Mo (ppm)
U(ppm) G
202
Hg (‰)
'
199
Hg (‰)
'
201
Hg (‰)
'
200
Hg (‰)
E1
E2
E3
c. TY
III
II
I
0 1000 2000
0 3000 6000
0.0 0.4 0.8
0.0 0.4 0.8
0 1000
01000
020
020
-3 -2 -1 0.1 0.2 0.3 -0.1 0.0 0.1
Hg (ppb)
Hg/TOC
Fe
HR
/Fe
T
Fe
PY
/Fe
HR
Hg/S
py
Hg/Al
Mo (ppm)
U(ppm)
G
202
Hg (‰) '
201
Hg (‰)
'
199
Hg (‰)
'
200
Hg (‰)
E1
d. WA
Glacial
diamictite
Siltstone/silty
mudstone
Siliceous
shale
Mostly
covered
Siltstone/silty
mudstone
Dolostone
Silty/standy
dolostone
Chert
Shaly
dolostone
Olistostrome
carbonate
Phosphorite/
chert
Dolomitic
phosphorite
Ooids/peloids/
intraclasts
Black
shale
Phosphorite
nodules
Bedded
pyrites
Fine-grained
sandstone
Exposure/dissolution/
erosional surface
Regional discontinuity
(abrupt facies change)
Fig. 2 | Chemostratigraphy of all study sections of the Doushantuo Formation
in South China. a Wuhe (WH), bYuanjia (YJ), cTaoying (TY) and dWeng’an (WA)
sections. Fe speciation,redox sensitiveelements (U, Mo),δ34S
pyrite
,Pconcentration
and P/Al are from ref. 8.“TOC”stands for total organic carbon. The horizontal blue
bands mark the intervals of previously proposed ocean oxygenation event
(corresponding to the E intervals). The gray band in the Δ199Hg and Δ201Hg column
of panel arepresents the analytical uncertainty of Hg odd-mass independent
fractionation (±0.04‰). The error bars for Hg isotopes are analytical uncertainties
defined in the method section.
Article https://doi.org/10.1038/s41467-023-39427-z
Nature Communications | (2023) 14:3920 3
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corresponding intervals in other sections as “E1”,“E2”and “E3”,
respectively (“E”standards for excursion). In WH, E1 shows the stron-
gest negative excursion (with Δ199Hg decreasing by as much as−0.29‰
relative to the highest value between the E intervals), whereas both E2
and E3 show smaller but still notable negative excursions (with Δ199Hg
dropping by −0.19‰). The other deep-water section, YJ, also shows
negative Δ199Hg values (down to −0.15‰) at the bottom of E1, which
then gradually increase to 0.01‰up section. In contrast, the two
shallower sections, WA and TY, show mostly positive Δ199Hg in the E
intervals (up to 0.24‰and 0.19‰, respectively), which is similar to the
positive values in WH between the E intervals.
The δ202Hg values (representing mass dependent fractionation,
MDF) of WH show a strong negative correlation with Δ199Hg (the
Δ199Hg/δ202Hg slope = −0.09 ± 0.02, R2=0.47, P<0.001), withpositive
excursions (up to 0.09‰) concurrent with negative shifts in Δ199Hg,
and gradual negative shifts (down to −2.60‰) concurrent with positive
shifts in Δ199Hg (Figs. 2and 3). The WA, TY and YJ sections also show
negative correlations between Δ199Hg and δ202Hg with similar Δ199Hg/
δ202Hg slopes as in WH (Supplementary Fig. S2). Collectively, all four
sections reveal a coherent linear Δ199Hg/δ202Hg correlation with a slope
of −0.10 ± 0.01 (R2= 0.62, P< 0.001) (Fig. 3a).
In addition to Δ199Hg/δ202Hg, the relationships between different
MIF values (Δ200Hg/Δ199Hg and Δ199Hg/Δ201Hg) are also diagnostic of the
mechanism of Hg isotope fractionation. The variation of Δ200Hg
(representing MIF of even isotopes) in each single section is within the
analytical uncertainty, but when all sections are combined, significant
positive relationships between Δ200Hg and Δ199Hg, and hence negative
correlations between Δ200Hg and δ202Hg can be observed, with a
Δ200Hg/Δ199Hg slope of 0.10 ± 0.01 (1SE, R2=0.55, P<0.001) (Fig. 3b)
and a Δ200Hg/δ202Hg slope of −0.01 ± 0.00 (1SE, R2=0.46, P< 0.001)
(Supplementary Fig. S3). Δ199Hg and Δ201Hg also show a significant
positive relationship with a Δ199Hg/Δ201Hg slope of 1.49 ± 0.09
(R2= 0.96, P< 0.001) for all sections (Supplementary Fig. S4).
The influence of Hg sources on Hg isotope compositions in
Doushantuo shales
The Hg isotope compositions in the study sections are unlikely mod-
ified by diagenetic or metamorphic alterations (see a detailed discus-
sion in Supplementary Text S2), and thus their variations should
represent changes in the primary Hg isotope signals of con-
temporaneous seawater, which could be affected by Hg sources or
fractionation processes in ocean. It is important to consider all pat-
terns of Hg isotope fractionation (including the direction and extent of
the shifts in both MIF and MDF values as well as their relationships),
because together they provide “multi-dimensional”constraints on the
mechanism of Hg isotope fractionation and thus allow us to constrain
the processes driving these fractionations. Since WH comprises a
continuous deposition of black shales from Member II to IV, most
discussion below focuses on WH.
The main Hg sources to the modern ocean include atmospheric
deposition and terrestrial input21, which have distinguishable Hg iso-
tope signatures (Supplementary Fig. S5 and Supplementary Text S3).
Enhanced input of terrestrial Hg was often invoked to interpret nega-
tive shifts in Δ199Hg of sedimentary rocks21, based on the fact that
modern terrestrial Hg associated with organic matter (OM) derived
from soil and biomass typically exhibits negative Δ199Hg27. However,
enhanced terrestrial Hg input is unlikely the main mechanism for the
negative Δ199Hg shift in the E intervals. Precambrian terrestrial biomass
was very scarce compared to today, except fungus-like microorganism
found in cryptic karstic environments28, and thus the modern-type
OM-associated terrestrial Hg was unlikely a major source of Hg. Before
the emergence of land plants, the dominant form of terrestrial Hg was
likely geogenic Hg released by weathering and erosion of continental
rocks, which is characterized by near-zero MIF29 and thus could not
explain the significantly negative Δ199Hg values in E1 of WH and YJ
sections. Although the near-zero Δ199Hg in E2 and E3 of WH seem to be
consistent with that of geogenic Hg, the correlations between THg and
-3.0 -2.5 -2.0 -1.5 -1.0 -0.5 0.0 0.5 1.0
-0.4
-0.3
-0.2
-0.1
0.0
0.1
0.2
0.3
0.4
0.5
0.6
-0.3 -0.2 -0.1 0.0 0.1 0.2 0.3 0.4 0.5
-0.10
-0.05
0.00
0.05
0.10
0.15
Modern atmospheric Hg(0)
Modern atmospheric Hg(II)
Modern open ocean sediment
Modern terrestrial Hg
Direct volanic emission
Stronger PZE
Late Devonian PZE
Mesoproterozoic PZE
TY (this study)
YJ (this study)
WA (this study)
WH (this study)
Δ
199
Hg ( )
δ
202
Hg ( )
a
Δ
200
Hg ( )
Δ
199
Hg ( )
b
StrongerPZE
Fig. 3 | Stable isotope compositions of Hg in all study sections. a Δ199Hg vs.
δ202Hg and bΔ200Hg vs. Δ199Hg for Weng’an (WA), Taoying (TY), Wuhe (WH),
and Yuanjia (YJ) sections. Common Hg sources based on modern samples
(average ±1 SD) are also plotted to show their po ssible contributions to Hg
isotope variations of the Doushantuo samples. See Supplementary Fig. S5 and
Supplementary Text S3 for details and references for modern data. The “Meso-
proterozoic PZE”and “Late Devonian PZE”are the Hg isotope data in sedimen-
tary rocks deposited under photic zone euxinia (PZ E) from the Mesoproterozoic
Atar and El Mreiti groups16, and from the Late Devonian Chattanooga Shale39,
respectively. The black short-dash lines are the linear regressions for all
Doushantuo samples (in panel a, slope = −0.10 ± 0.01, R2= 0.62, P< 0.001; in
panel b, 0.10 ± 0.01, R2= 0.55, P< 0.001). The red dash lines are reference lines
of experimental results of dark abiotic oxidation of Hg(0) by thiol compounds
(in panel a, slope = −0.12; in panel b, slope = 0.13)46. The error bars are analytical
uncertainties defined in the method section.
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Nature Communications | (2023) 14:3920 4
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Al (a proxy of continental weathering input) are insignificant in the
entire WH section as well as in YJ and two E intervals of TY sections
(Supplementary Text S1 and Supplementary Fig. S1), which strongly
argue against enhanced terrestrial Hg input as the cause of the nega-
tive shift in Δ199Hg during E intervals. In addition, the Hg concentration
in crustal igneous and metamorphic rocks is typically <10ppb30,31,
much lower than the >100ppb (and frequently >1000ppb) THg in
almost all Doushantuo shales, which also argues against terrestrial Hg
from continental weathering as a dominant source of Hg. Moreover,
we simulated the shift in Δ199Hg caused by enhanced terrestrial
weathering using a simplified Hg isotope box-model (see details of the
model in Supplementary Text S4)32, and found that even under a sce-
nario of 50× increase in terrestrial Hg input, the Δ199Hg of the ocea n can
only be decreased by 0.02‰(Fig. 4),whichisanorderofmagnitude
lower than the extent of Δ199Hg excursions during the E intervals
(−0.19‰to −0.29‰).ThereasonforthemutedΔ199Hg shift under high
terrestrial inputs is that the enhanced terrestrial input would increase
the sizeof the oceanic Hg reservoir, whichwould also lead toenhanced
re-emission of Hg from ocean to the atmosphere as observed in
modern ocean33,34. The re-emitted Hg would undergo atmospheric
redox transformations that produce a net positive Δ199Hg signal for
atmospheric Hg species35, which then deposit back to surface ocean.
Thus, the enhanced re-emission and re-deposition of Hg with positive
Δ199Hg following enhanced terrestrial weathering eventually counter-
acted the negative shift in Δ199Hg caused by terrestrial input. A full
investigation on the effects of Hg re-emission and re-deposition on
marine Hg isotope signatures is beyond the scope of the current study
and is being undertaken in a separate study. Nevertheless, the mod-
eling results further strengthen our argument that terrestrial Hg could
not account for the negative Δ199Hg shift during E intervals.
Atmospheric Hg deposition is the dominant Hg source to global
ocean in modern environment36, and thus was likely a major Hg source
in the Ediacaran ocean as well. The Hg isotope compositions of mod-
ern open ocean seawater are proposed to reflect mixing between
atmospheric Hg(II) and Hg(0) deposition36, which have opposite iso-
tope signatures, with positive Δ199Hg, positive Δ200Hg and negative
δ202Hg for the former, and negative Δ199Hg, negative Δ200Hg and posi-
tive δ202Hg for the latter (Supplementary Fig. S5 and reference therein).
Therefore, a simple explanation for the cyclic variation of Hg isotopes
in WH is a change in the proportion of atmospheric depositions, i.e.,
with a higher proportion of Hg(0) deposition during intervals with
relatively more negative MIF (the E intervals) and a higher proportion
of Hg(II) deposition during intervals with more positive MIF (between
the E intervals). However, this simple interpretation is inconsistent
with several observations. First, the Δ199Hg/Δ201Hg of both atmospheric
Hg(II) and Hg(0) is ~1.0 (typically between 1.0 and 1.1)29, whereas the
Δ199Hg/Δ201Hg of Doushantuo shales is ~1.5 (Supplementary Fig. S4),
which cannot be solely explained by the mixing of atmospheric Hg(II)
and Hg(0). Second, changes in the proportions of atmospheric
deposition to modern ocean have been shownto cause corresponding
variation of Δ200Hg by more than 0.1‰in marine sediments36, because
of the opposite Δ200Hg for atmospheric Hg(0) and Hg(II) (Supple-
mentary Fig. S5), but this is inconsistent with the invariant Δ200Hg in
WH (Fig. 2).
In addition to terrestrial input and atmospheric deposition,
upwelling of Hg associated with OM (Hg-OM) from the deep ocean has
also been proposed to drive the Hg isotope variation in the Doush-
antuo Formation by a previous study on coeval shallow shelf lagoon
sections (e.g., Jiulongwan section)37. This study found primarily posi-
tive Δ199Hg shifts and negative δ202Hg shifts that are roughly in sync
0
5
10
15
20
640 630 580 570 560
0.00
0.04
0.08
0.12
0.16
0
1000
2000
3000
4000
5000
10 x weathering flux
20 x weathering flux
50 x weathering flux
Weathering flux (Mg a
-1
)
aE1 E2 E3
0
200
400
600
800
1000
1200
Local volcanic flux (Mg a
-1
)
Enrichment factor_ocean
b
'
199
Ma
c
Fig. 4 | Numerical simulation results using a simplified Hg isotope box-model.
Simulated marine Hg enrichment (b)andΔ199Hg (c) in response to an increase in
terrestrial weathering Hg flux by 10–50× during E1 (~635 Ma), E2 (~580 Ma), and E3
(~565Ma), and a further increase in local volcanic Hg emission by 100× during E3
(a). The “enrichment factor_ocean”value is defined as the marine Hg reservoir
during E intervals relative to the reservoir size between E intervals.
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Nature Communications | (2023) 14:3920 5
Content courtesy of Springer Nature, terms of use apply. Rights reserved
with the negative carbonate carbon isotope (δ13C
carb
) excursion in the
cap dolostone (Member I) and Member IV in shelf lagoon sections.
These Hg isotope signals were proposed to reflect the Hg released
during the oxidation of deep ocean OM that is enriched in light carbon
isotopes37. The isotopic signatures of Hg associated with deep ocean
OM are hypothesized to resemble open ocean sediments38.Sincethe
oxidation of OM is oxygen-demanding, this source should be more
prominent during the proposed OOE intervals and in shallow water
sections where the upwelling DOC meets the chemocline and is oxi-
dized. In our study sections, Hg-OM is unlikely a dominant source for
deep-water sections (WH and YJ), because they actually show opposite
shifts in Δ199Hg and δ202Hg during the E intervals than the purported
Hg-OM isotope signature (Fig. 2). E1 of the two shallower sections (WA
and TY) shows primarily positive Δ199Hg that is similar to modern open
ocean sediments, and strong correlations between THg and TOC
(Supplementary Text S1 and Supplementary Fig. S1), indicating possi-
ble contributions of Hg-OM. However, Hg-OM alone still cannot
explain the high Δ199Hg/Δ201Hg slope for all sections (~1.5) (Supple-
mentaryFig. S4), which is significantly higher than those of open ocean
sediments (close to 1.0)38.
Therefore, while terrestrial Hg, atmospheric Hg, and deep ocean
Hg-OM may all constitute the sources of Hg in the Doushantuo shales,
the variations of Hg isotopes could not be fully explained by changes in
or mixing of these Hg sources, but require fractionation of Hg isotopes
in ocean.
Hg isotope evidence for PZE in the Ediacaran ocean
The distinct negative excursions of Δ199Hg with concurrent positive
excursions of δ202Hg during the E intervals were most likely driven by
Hg isotope fractionation in seawater under PZE conditions, and this
hypothesis is supported by abundant and compelling evidence. First,
there exists a strong positive correlation between Δ199Hg and pyrite S
isotopes (δ34S
pyrite
)inWHsection(R
2=0.83, P<0.001) (Fig. 5). Very
negative δ34S
pyrite
values (as low as −34.9‰) are found during the
proposed OOE intervals at WH and have been attributed to an
increased availability of marine sulfate, which permits the expression
of large-magnitudes of S isotope fractionation during bacterial sulfate
reduction7. Increased marine sulfate availability would also increase
H
2
S production, fueling the expansion of euxinia. Thus, the highly
synchronized shifts in Δ199Hg and δ34S
pyrite
suggest a first-order control
of redox conditions on the variation of Hg isotopes. Furthermore, the
THg of WH shows the strongest correlation with S
py
among all host
phases in at least two E intervals (corresponding to Member II and
Member IV OOEs) (Supplementary Text S1 and Supplementary Fig. S1),
suggesting that Hg was primarily scavenged to sediments by sulfide
minerals. The strong correlation between THg and sulfide is typically
interpreted as a sign of euxinic depositional environment26.Incon-
trast, between the E intervals, THg shows a stronger correlation with
TOC than with S
py
(Supplementary Text S1 and Supplementary Fig. S1),
suggesting that the sedimentation of Hg was controlled by OM, like in
modern open oceans21.Thus,thechangeinhostphasesfromsulfide to
OM likely indicates a transition from locally euxinic to non- or less
euxinic conditions.
Second, the patterns of Hg isotope excursions (including the
direction and extent of the shifts in both Δ199Hg and δ202Hg values as
well as Δ199Hg/δ202Hg) in the E intervals are strikingly similar to the Hg
isotope systematics from two previously reported successions where
PZE was validated independently. Zheng et al. reported Hg isotopic
data from Mesoproterozoic black shales (~1.1 Ga) of the Atar and El
Mreiti groups, West Africa16,andfromtheLateDevonianChattanooga
Shale, North America39, and both display comparable negative shifts in
Δ199Hg (by as much as ~ −0.2‰) during PZE with similar negative
Δ199Hg/δ202Hg of −0.15 and −0.10, respectively (Fig. 3). The PZE con-
dition in the two above-mentioned studies was independently vali-
dated by organic biomarkers produced by anoxygenic phototrophic
bacteria that metabolize H
2
S40,41. While indigenous biomarkers are not
yet available for the Doushantuo shales42, the highly similar Hg isotope
signatures between these successions serves as compelling evidence
of a shared driving mechanism.
As proposed previously, PZE can result in concurrent negative
Δ199Hg shift and positive δ202Hg shift via two mechanisms16.Thefirst is
photoreduction of Hg(II) bound to reduced organic sulfur ligands [i.e.,
Hg(II)-S] in a sulfide-rich photic zone. Under sulfidic conditions, dis-
solved organic matter in seawater will likely have a high fraction of
thiols due to sulfurization43, favoring the formation of Hg(II)-S com-
plexes. Multiple experimental studies have demonstrated that the
photoreduction of Hg(II)-S increases δ202Hg and decreases Δ199Hg in
the remaining Hg(II) phase44,45, which is consistent with the pattern
of Hg isotope shifts observed in the E intervals. One caveat is that
the experimental Δ199Hg/δ202Hg slope of this photoreduction process
(~−0.7) is different from that of the Doushantuo shales (−0.10)44,
suggesting that the photoreduction of Hg(II)-S is not the only
process involved.
The second mechanism is enhanced oxidation of atmospheric
Hg(0) in seawater induced by the abundant thiol compounds under
PZE conditions. This mechanism is fully compatible with the Hg iso-
tope signatures of Doushantuo shales. Our previous experimental
study shows that thiol-induced aqueous Hg(0) oxidation leads to
negative shifts in Δ199Hg (by as much as −0.20‰), positive shifts in
δ202Hg and no changes in Δ200Hg in the oxidized Hg(II) phase, with a
characteristic Δ199Hg/δ202Hg of −0.12 and a Δ199Hg/Δ201Hg of 1.646,which
are all fully consistent with the pattern of Hg isotope variation in the E
intervals (Fig. 3and Supplementary Fig. S2). Although it may seem
counter-intuitive that Hg(0) oxidation is enhanced in sulfidic water
that is typically considered as a reducing environment, multiple
experimental studies indeed show that aqueous Hg(0) oxidation can
be enhanced by either free thiol or thiol ligands in organic matter46,47.
Sulfide also inhibits the re-emission of Hg(0) from seawaters by com-
plexing and scavenging the oxidized Hg to sediments. Thus, the pre-
sence of PZE could promote the marine uptake of atmospheric Hg(0)
by enhancing its oxidation in the surface ocean and subsequent
sequestration to sediments. Moreover, the rate of Hg(0) oxidation in
sulfidic waters determined by experiments is comparable to many
-0.1 0.0 0.1 0.2
-40
-20
0
20
Δ199Hg ( )
δ34Spyrite ( )
R2=0.83, P< 0.001
Fig. 5 | Positive correlation between Δ199Hg and δ34S
pyrite
in Wuhe section. The
black solid line represents the linear regression between Δ199Hg and δ34S
pyrite
(R2=0.83, P<0.001).
Article https://doi.org/10.1038/s41467-023-39427-z
Nature Communications | (2023) 14:3920 6
Content courtesy of Springer Nature, terms of use apply. Rights reserved
other major redox pathways of Hg in natural aquatic systems (e.g.,
photooxidation and photoreduction)47, making the thiol-induced
Hg(0) oxidation an environmentally relevant process and potentially
the key mechanism that caused the observed shifts in Hg isotopes
during periods of PZE.
In contrast to the E intervals, the gradual positive shifts in Δ199Hg
with concurrent negative shifts in δ202Hg between the E intervals
indicate the contraction of PZE and the transition to modern-like
oxygenated surface ocean, which is supported by the change in Hg
host phases from sulfide to OM (Supplementary Fig. S1), as well as the
similarly positive Δ199Hg and negative δ202Hg in modern ocean
sediments36,38,48 (Supplementary Text S3 and Supplementary Fig. S5).
Note that this interpretation does not conflict with the Fe speciation
data between the E intervals (Fe
HR
/Fe
T
>0.38,Fe
py
/Fe
HR
>0.8)(Fig.2),
which is typically interpreted to reflect local bottom water euxinia49.
Sahoo et al. who first reported these Fe speciation data suggested that
the high Fe
py
/Fe
HR
could potentially reflect sulfidic conditions in
sediment porewater rather than the water column7.Incontrast,Hg
isotopes specifically track euxinia in the photic zone, because the two
proposed mechanisms controlling Hg isotope fractionation under PZE
both operate in surface water where light can penetrate or atmo-
spheric gaseous Hg(0) can dissolve. There is currently no known
mechanism that produces characteristic Hg isotope signatures in
euxinic (but non-PZE) bottom water or sediment porewater, where
sulfide only acts as a ligand that binds with Hg. In fact, the study by
Zheng et al.16 found that the Mesoproterozoic deeper-water section
deposited under euxinic (but not necessarily PZE) conditions shows no
or a much weaker negative shift in Δ199Hg than the shallower section
wherePZEoccurred,suggestingthatHgisotopestrackspecifically PZE
rather than euxinia in general.
Spatial variation of PZE in the South China Basin
The Hg isotope chemostratigraphic data are incomplete for three of
the four study sections (YJ, TY, and WA) because they are not com-
pletely preserved, thus we are unable to perform a “full profile”
reconstruction for redox conditions in these sections as for WH.
However, all four sections have Hg isotope data for E1 and they show
clear variations of Hg isotopes within E1. By comparing the variation
patterns of Hg isotopes during E1 across the four sections, which cover
a transect from the shelfto the basin, we are able to evaluate the spatial
extent of PZE and the dynamics of the euxinic water mass within the
Ediacaran South China Basin. The E1 of WH also shows the strongest
negative shift in Δ199Hg among all E intervals, possibly indicating the
most extensive PZE event. Thus, our discussion below focuses on E1.
Among the four study sections, the basinal section (YJ) shows the
most negative Δ199Hg and most positive δ202Hg at the lowermost E1,
and it also exhibits an up-section positive shift in Δ199Hg and con-
current negative shift in δ202Hg with a Δ199Hg/δ202Hg ratio similar to
that in WH, suggesting that PZE was likely the most extensive at the
onset of E1 but weakened towards the end of E1 at YJ (Fig. 2). In con-
trast, both the lower-slope (WH) and upper-slope (TY) sections show
clear up-section negative shifts in Δ199Hg and concurrent positive shifts
in δ202Hg with similar Δ199Hg/δ202Hg ratios, suggesting a gradual
expansion of PZE at these two sites. The WA section located on the
shallow shelf margin has the most positive Δ199Hg and the most
negative δ202Hg among the four sections, and its THg correlates with
TOCandAl,butnotwithsulfide (Supplementary Text S1 and Supple-
mentary Fig. S1), suggesting that WA was likely located above the
euxinic water mass or the PZE at WA was too sporadic to be captured
by the current data. This is in agreement with previous paleogeo-
graphic reconstruction of the Doushantuo Formation in South China,
which suggests that WA was likely exposed to oxic/suboxic water
above the chemocline22.
Based on the above observations, we can reasonably propose that
PZE was widespread during the deposition of the basal Doushantuo
shale in South China, extending at least from the upper slope to the
basin, which is consistent with the previous proposal for a mid-depth
euxinic wedge on continental shelves of the South China Basin11,12.
Furthermore, the euxinic wedge probably became shallower
approaching the end of E1, and this is evident from the different pat-
terns of Hg isotope variations within E1 across different sections. As
discussed above, PZE was weakened at the deepest YJ section towards
the end of E1, while PZE intensified at the shallower WH and TY sec-
tions, suggesting an up-slope migration of the euxinic water mass
toward the end of E1 as depicted by Fig. 6. It is worth noting that E1
coincided with a sea level highstand23. Thus,the up-slope migration of
theeuxinicwedgetowardstheendofE1couldbeassociatedwitha
global or local sea level change, although this hypothesis needs to be
tested further.
Plausible drivers of PZE
The occurrence of PZE during the previously proposed OOE intervals
was likely a result of increased marine sulfate and nutrient availability,
which has been previously proposed to trigger the Member II OOE
(~635 Ma, E1) due to enhanced continentalweathering in the aftermath
of the Marinoan glaciation7. During the process of deglaciation, it is
hypothesized that a surge of nutrients from freshly exposed land
surfaces and finely ground glacial tills was delivered to the ocean,
which increased primary productivity and thus oxygen consumption
by newly produced organic matter, leading to anoxia on a timescale of
~104years2,50. The enhanced primary productivity would also increase
the organic matter available for microbial sulfate reduction and hence
seawater H
2
S accumulation in near-shore environments, leading to
PZE. On a longer times cale (~106years), increased nutrient influx would
increase organic carbon burial and hence atmospheric oxygen50.A
similar sequence-of-events may have followed the Gaskiers glaciation,
which occurred broadly coeval with the Member III OOE (~580 Ma,
E2)8. However, this link is speculative given the uncertain timing of the
Member III OOE, and it merits further examination.
The increased availability of nutrients in global oceans after the
Marinoan glaciation is well-documented by the apparent expansion of
themarinephosphorous(P)reservoir
51–54. Phosphorous is a key limit-
ing nutrient and increased P influxes to the ocean have been proposed
to account for the development of PZE in many cases across Earth
history, particularly during various Phanerozoic mass extinction
events55,56, and inmodern meromictic lakes (e.g., the Mahoney Lake)57.
Both P and P/Al (to account for changes in sedimentation rate) in the
WH section show a cyclic variation similar to Hg isotopes and other
redox proxies, with lower P during the E intervals and higher P in
between (Fig. 2). The accumulation of P in sediments is controlled by
local redox conditions, and euxinic conditions would re-dissolve and
recycleP in sedimentsback to the watercolumn due to rapid reduction
of P-bearing Fe oxides by H
2
S and preferential release of P from
remineralization of organic matter by sulfate-reducing bacteria58.
Thus, the variation of sedimentary P is consistent with the local redox
conditions at WH, and suggests intensive recycling and regeneration of
P modulated by a rising marine sulfate reservoir54, which may have
stimulated PZE via positive feedback.
The Member IV OOE (E3) of both WH and TY sections shows
higher Hg enrichments and different patterns of Hg isotopes com-
pared to other OOEs, which we ascribe to Hg input from local volcanic
emission (for WH) or upwelling of deep ocean Hg-OM (for TY) (see
detailed discussion in Supplementary Text S5). Despite the influences
of volcanic or upwelling Hg, the E3 interval of WH still shows a strong
correlation between THg and S
py
(Supplementary Fig. S1). Further-
more,wealsosimulatedtheshiftinΔ199Hg caused by enhanced local
volcanic Hg input using the Hg isotope box-model (more details in
Supplementary Text S4), and found that an increase in local volcanic
Hg flux by 100× could only decrease Δ199Hg by <0.03‰(Fig. 4), which
is similar to the muted Δ199Hg shift caused by terrestrial Hg input. As
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Nature Communications | (2023) 14:3920 7
Content courtesy of Springer Nature, terms of use apply. Rights reserved
explained earlier, the reason for the muted Δ199Hg shift might be due to
the re-emission of volcanically-sourced Hg from ocean and the re-
deposition of atmospheric Hg with positive Δ199Hg back to ocean.
Thus, the negative excursion E3 with a drop in Δ199Hg by −0.19‰may
also suggest PZE. This hypothesis is consistent with the previous
proposal for widespread euxinia on continental shelves based on V
isotopes in black shales of Member IV59. Volcanism and upwelling may
also lead to enhanced nutrient availability, which may trigger PZE,
although this hypothesis requires further verification.
The negative impacts of PZE on ocean oxygenation and the
Ediacara biota
PZE may have exerted a negative impact on the oxygenation of the
Ediacaran ocean by promoting anoxygenic photosynthesis and inhi-
biting oxygenic photosynthesis (Fig. 6). This may be an
underappreciated mechanismfor the transient nature of the Ediacaran
OOEs. While oxygenic photosynthesis is the overwhelming source of
free O
2
, it can be outcompeted by anoxygenic photosynthesis under
conditions that favor the latter, such as extensive PZE60,61 or ferrugi-
nous conditions62. Compelling experimental evidence has shown that
even μMlevelsofH
2
S can completely shut down oxygenic
photosynthesis63. It has been suggested that the competition for light
and nutrients between anoxygenic and oxygenic photosynthetic
organisms was a key mechanism responsible for the delayed oxyge-
nation of Earth’satmosphere
62, and for the relatively low level of
oxygen throughout the mid-Proterozoic18. Although our Hg isotope
data only cover South China, lipid biomarkers for PZE have been
reported in late-Ediacaran petroleum from Eastern Siberia64,suggest-
ing that PZE may have been common in various continental margins
during the Ediacaran Period. Continental margins are the major
Photic zone
YJ
WH
TY
WA
3. Expanded and dynamic PZE limited
oxygen production and habitable space
2. P regeneration
enhanced PZE
1. Enhanced inputs of
continental nutrients and
sulfate triggered PZE
Oxic/suboxic Ferruginous Euxinic PZE
a. Intervals of expanded PZE
b. Intervals of contracted PZE
Photic zone
YJ
WH
TY
WA
3. Contracted PZE
2. Inhibited P
regeneration
1. Weathering inputs
waned down 4. Decreased photoreduction of Hg(II)-S and
oxidation of atmospheric Hg(0) in photic zone
biota
No fossil
No fossil
Wenghui biota
in upper
Doushantuo
biota
No fossil
No fossil
Wenghui biota
in upper
Doushantuo
+
-
+-0
0
199
Hg
202
Hg
+
-
+-0
0
199
Hg
202
Hg
H
2
S
H
2
S
Hg(0) Hg(II)-S
4. PZE promoted photoreduction of Hg(II)-S and
oxidation of atmospheric Hg(0) in photic zone
Fig. 6 | Photic zone euxinia (PZE) in the Ediacaran ocean and its impact on
ocean oxygenation and the Ediacara biota. Conceptual shifts in Δ199Hg and
δ202Hg under the depicted redox conditions are shown in inserted boxes. The
numbers in text boxes represent the sequence of events. During intervals of
expanded PZE (panel a), PZE developed on continental margins due to increased
marine nutrients and sulfate availability (likely due to enhanced postglacial con-
tinental weathering, particularly during E1), and was further promoted by the
positive feedback from P regeneration under euxinic water. PZE would have
enhanced the photoreduction of Hg(II)-S (“S”represents reduced sulfur ligands)
and the oxidation ofatmosphericallydeposited Hg(0), leading to the negative shift
in Δ199Hg with concurrent positive shift in δ202Hg. Expanded PZE also promoted
anoxygenic photosynthesis and inhibited oxygenic photosynthesis, resulting in
relatively rapid termination of the previously proposed ocean oxygenation event.
Spatial variations in Hg isotopes indicate that PZE may have migrated upslope
toward the end of E1, as depicted by the upper euxinic wedge in panel a.During
intervals of contracted PZE (panelb), the postglacial surge ofweathering may have
waned down, leading to the contractionof PZE and resultingin the negative shift in
δ202Hg with concurrent positive shift in Δ199Hg. PZE may have also limited the
diversification of complex eukaryotic life. At Weng’an (WA) section that was not
affected by PZE, the Weng’an biotawas found in the upper phosphorite.At Taoying
(TY) section where PZE was likely sporadic, the Wenghui biota was found in the
upper Doushantuoblack shales. In contrast, at Wuhe (WH)and Yuanjia (YJ) sections
where PZE or anoxic water persisted, no benthic fossils have been reported.
Article https://doi.org/10.1038/s41467-023-39427-z
Nature Communications | (2023) 14:3920 8
Content courtesy of Springer Nature, terms of use apply. Rights reserved
habitats for early primary producers such as cyanobacteria and
ancestral photosynthetic eukaryotes61. Thus, we propose that the
widespread PZE in continental margins initiated a negative feedback
on oxygenic photosynthesis, which may be partially responsible for
the relatively rapid termination of OOE. This argument is consistent
with a recent hypothesis based on theoretical models that found the
cyclic and oscillating oxygenation and deoxygenation events during
the Ediacaran Period can be explained by internal feedbacks in the
biogeochemical cycles of carbon, oxygen and phosphorous, without
the need for specific external forcing65. As an important step forward,
our findingof PZE providesan independent validation and mechanistic
insights into this theoretical hypothesis.
Recurrent and widespread PZE in the Ediacaran ocean may have
also inhibited the diversification and ecological expansion of early
macroscopic eukaryotic life by limiting their habitable space via H
2
S
poisoning. H
2
Stoxification during PZE was a potent kill mechanism for
some of the largest mass extinctions, such as the end-Permian20,32,66
and end-Triassic mass extinctions67. Emerging geochemical data sug-
gest a redox control on the rise and fall of the Ediacara biota4,5,68.The
unstable redox conditions and repeated development of PZE in early
Ediacaran oceans as revealed by Hg isotopes likely limited the ecolo-
gical expansion of macroscopic multicellular eukaryotes. For example,
the lower slope and basinal facies, as represented by the WH and YJ
sections, were strongly influenced by PZE during the OOEs (the E
intervals) and were consistently anoxic between OOEs, excluding the
establishment of benthic multicellular eukaryotes that require free
oxygen. In relatively shallow-water shelf and upper slope facies, as
represented by the WA and TY sections, eukaryotes were able to
colonize the benthic realm where and when PZE was absent. For
example, multicellular eukaryote fossils of the Weng’an biota at WA
occur in phosphorite strata correlated with Member II above the E1
interval69. Similarly, macroalgal fossils of the Wenghui biota at TY
occur in black shales correlated to the Member III above the E2 interval
(Fig. 2). These fossiliferous intervals represent locally oxic environ-
ments (as indicated by negative Ce anomalies in fossiliferous
phosphorite70) that were habitable for early Ediacaran eukaryotes. A
limitation of this study is that we did not directly measure samples
from fossiliferous intervals. Thus, future studies that compare fossili-
ferous and non-fossiliferous intervals are needed to further verify the
proposed impact of PZE on the Ediacara biota.
In summary, our study reveals recurrent, widespread and spatially
dynamic PZE on the continental margin of South China during pre-
viously proposed OOE intervals of the Ediacaran Period. Although the
presence of a dynamic euxinic wedge at South China during the
Ediacaran Period has been proposed previously, our study shows
strong evidence for the invasion ofeuxinic water mass into the marine
photic zone, which has important implications for understanding its
impacts on the redox evolution and biological innovation in the
Ediacaran ocean. We suggest that widespread PZE at continental
margins was a consequence of increased marine nutrient and sulfate
reservoirs due to their positive feedback on primary productivity, but
PZE may have also initiated a series of negative feedbacks that inhib-
ited oxygen production (and thus decelerated ocean oxygenation) by
promoting anoxygenic photosynthesis and limiting the habitable
space foreukaryotes carrying out oxygenic photosynthesis. Therefore,
PZE could be one of the key reasons for the transient nature of ocean
oxygenation during the Ediacaran Period and the delayed rise of the
Ediacara biota.
Methods
Study sections
The locations, stratigraphy, and paleogeography of the study sections
are described in detail in previous work7,8,22, and thus only a brief
overview is presented here. The Doushantuo Formation is best known
in the Yangtze Gorges area where it is divided into four lithostrati-
graphic members. Member I serves as a regional stratigraphic marker
bed and corresponds to the cap carbonate overlying the terminal
Cryogenian-age glacial diamictite of the Nantuo Formation. A zircon
U-Pb date from a volcaniclastic layer atop Member I constrains its
deposition age to be 635 ± 0.6 Ma71. Member II consists primarily of
black shales with subordinate carbonate beds. Member III is domi-
nated by carbonate although black shale becomes increasingly
abundant toward deep-water facies outside the Yangtze Gorges area.
The Member II-III boundary is correlated with a regional stratigraphic
discontinuity in shallow-water facies22, and if this regional dis-
continuity tracks a widespread marine regression, it may be time-
equivalent with the Gaskiers glaciation (~580 Ma)8. Member IV is
dominated by organic-rich black shales, and can also be used as a
regional stratigraphic marker bed for the Doushantuo Formation in
South China. The age of Member IV is a topic of ongoing debate, but
it is probably ~565 Ma72. The shelf and upper slope sections (WA and
TY) contain Ediacaran microfossils of the Weng’an biota14 and mac-
rofossils of the Wenghui biota73, respectively. No fossils have been
reported from the WH and YJ sections that represent lower slope and
basinal facies.
Hg concentration analysis
Mercury concentrations in solid samples were analyzed on a Lumex
R-915F Hg Analyzer at Tianjin University. Mercury in samples was
released as Hg(0) vapor by combustion at a temperature of ~750 °C,
and then measured by a cold vapor atomic absorption spectroscopy
(CV-AAS) with a detection limit lower than 0.5 ng/g. A certified refer-
ence material GBW07311 (GSD-11, freshwater sediment) was measured
repeatedly alongside samples to monitor analytical accuracy and
reproducibility. The GSD-11 yielded an average Hg concentration of
72.2 ± 5.8 ng/g (2 SD, n= 8), which is consistent with its certified value
72 ± 9 ng/g.
Hg isotope analysis
Mercury isotopes were analyzed using multi-collector inductively
coupled plasma mass spectrometry (MC-ICPMS, Neptune Plus,
Thermo Fisher Scientific) at School of Earth System Science, Tianjin
Universitybasedonpublishedmethods
74. Prior to isotopic analysis,Hg
in samples was first extracted by acid digestion and then purified by
ion-exchange chromatography (see details in Supplementary
Text S6)16. To monitor Hg yields and the accuracy of Hg isotope
measurement, procedural blanks and three standard reference mate-
rials (SRM), including GBW07405 (Yellow/red soil), NIST 2702 (Inor-
ganics in marine sediment), USGS SBC-1 (Brush Creek Shale), and the
NIST SRM 3133 Hg isotope standard, were processed alongside sam-
ples. Recoveries for all SRM were 100 ± 19% (2 SD, n=19) (Supple-
mentary Table S1). Mercury concentrations in procedural blanks were
typically less than 1% of the Hg present in samples.
Solutions eluted from the chromatographic procedure were
diluted to 1.5 ng/g of Hg using a matrix solution containing 5% HCl.
Thereafter, Hg in the diluted solutions was reduced by SnCl
2
(3%, w/v)
to gaseous Hg(0), which wasthen carried into the plasma of MC-ICPMS
by Hg-free argon gas. Simultaneously, thallium (Tl) aerosol (NIST SRM
997) generated by Aridus II desolvator was introduced together with
Hg(0) vapor into the plasma. Five Hg isotopes (198Hg, 199Hg, 200Hg,
201Hg, and 202Hg) and two Tl isotopes (203Tl, 205Tl) were simultaneously
measured via Faraday cups. Instrumental mass bias was corrected for
using a combination of internal calibration with measured 205Tl/203Tl
ratios and standard-sample-standard bracketing (relative to the NIST
SRM 3133 Hg standard). The bracketing standard was matched to
samples in terms of both matrix and Hg concentration (less than 10%
difference). On-peak zero corrections were applied to all measured Hg
masses. Mercury isotope compositions are reported using δnotation
Article https://doi.org/10.1038/s41467-023-39427-z
Nature Communications | (2023) 14:3920 9
Content courtesy of Springer Nature, terms of use apply. Rights reserved
defined by the following equation:
δxHgðmÞ=ðxHg=198HgÞsam ple
ðxHg=198HgÞstd
1
"#
× 1000 ð1Þ
where XHg is 199Hg, 200Hg, 201Hg, or 202Hg, and “std”represents the NIST
SRM 3133 standard. The MDF is reported as δ202Hg, and MIF is reported
as the capital delta notation (Δ) according to the following equation:
ΔxHgðmÞ=δxHg ðδ20 2Hg × βÞð2Þ
where x is the mass number of Hg isotope 199, 200, and 201. βis a
scaling constant used to calculate the theoretical kinetic MDF, and it is
0.2520, 0.5024, and 0.7520 for 199Hg, 200Hg, and 201Hg, respectively75.
To ensure data quality,each sample was measured at least twice,
and a commonly used reference standard NIST 8610 was measured
every 6 −7 samples to monitor instrument performance. The
averages of all NIST 8610 are: δ202Hg = −0.54 ± 0.06‰,
Δ199Hg = −0.03 ± 0.04‰,Δ200Hg = 0.00 ± 0.02‰(2 SD, n=39), con-
sistent with the published values75. The NIST SRM 2702 yielded
average δ202Hg, Δ199Hg and Δ200Hg values of −0.77 ± 0.04‰,
−0.03 ± 0.06‰, and 0.00 ± 0.02‰(2SE, n= 3), respectively (Supple-
mentary Table S1). All isotope data are reported in Supplementary
Table S2 and analytical uncertainties are reported as either 2 stan-
dard error (2SE) of sample replicates or 2 SD of all measurements of
the NIST 8610, whichever is higher.
Data availability
The authors declare that all data reported this study are available in the
Supplementary Information file and at Figshare https://doi.org/10.
6084/m9.figshare.23003246.
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Acknowledgements
This work was funded by the National Natural Science Foundation of
China (grant No. 41973009, 41830647), National Key R&D Program of
China (2022YFF0800300) and the National Science Foundation (EAR
2021207 awarded to S.X., and EAR 1760203 to A.D.A.).
Author contributions
W.Z., S.K.S. and J.C. conceived, designed and supervised the study. A.Z.
performed the analyses of Hg concentration and Hg isotopes. R.S.
performed the modeling work. W.Z., J.C., S.X. and A.D.A. acquired
funding for this study. W.Z. led the paper writing with significant inputs
Article https://doi.org/10.1038/s41467-023-39427-z
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from M.R.N., C.M.O., A.D.A., and S.X. All authors contributed to data
interpretation and writing.
Competing interests
The authors declare no competing interests.
Additional information
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