Introduction
The recovery of marine invertebrate faunas and ecosystems after the
∼ 252 Ma end-Permian mass extinction appears to have been the
most protracted following any Phanerozoic biocrisis (Erwin, 2001; Bottjer et
al., 2008). As with the mass extinction event, many aspects of the Early
Triassic recovery remain uncertain, including its timing, pattern, and
causes. Species origination rates and biodiversity did not return to
pre-extinction levels until the early Middle Triassic, after a protracted
process of niche building and increasing ecosystem complexity (Chen and
Benton, 2012). The slowness of the recovery process is believed to have
resulted, in part, from the effects of sustained or repeated environmental
stresses during the Early Triassic (Algeo et al., 2011; Retallack et al.,
2011). In particular, the pace of the biotic recovery may have been related
to episodic large-scale injection of volcanic CO2 and thermogenic
CH4 into the atmosphere, probably from the Siberian Traps large igneous
province, and a resulting intensification of ocean anoxia (Retallack and
Jahren, 2008; Black et al., 2012).
The extreme environmental conditions (tropical SSTs (sea-surface temperature)
> 35 ∘C) of the first ∼ 1.5 Myr of the Early
Triassic came to an end at the ∼ 250 Ma Smithian–Spathian boundary
(SSB), which subdivides the Olenekian stage of the Lower Triassic and which
is defined by the first appearance of the conodont Novispathodus pingdingshanensis at Chaohu, Anhui Province, eastern China (Zhao et al.,
2007). The SSB witnessed major changes among marine biotas, including a
severe loss of biodiversity among conodonts and ammonoids (Orchard, 2007;
Stanley, 2009; Brayard et al., 2009), size reduction (Lilliput effect) among
surviving conodont taxa (Chen et al., 2013), and a contraction of the
paleolatitudinal range of surviving ammonoid taxa (Galfetti et al., 2007;
Brayard et al., 2009). The SSB also marked a major change in global climate,
with strong tropical sea-surface cooling (Sun et al., 2012; Romano et al.,
2013) and a steepening of the latitudinal temperature gradient (Galfetti et
al., 2007). To date, however, the SSB event has received detailed study only
in several sections in South China (Galfetti et al., 2007; Liang et al.,
2011) and the Salt Range of Pakistan (Hermann et al., 2011). Here, we report
the SSB event from a new Lower Triassic section in southern Guizhou Province,
South China. We correlate this section with existing SSB sections using a
combination of conodont biostratigraphic and carbon isotopic constraints, and
we examine changes in marine environmental conditions using a combination of
elemental and isotopic proxies, with the goal of better understanding the
role of the SSB in the recovery of Early Triassic marine ecosystems.
Smithian–Spathian boundary at the study section
The study section (GPS: 25∘45′9.6′′ N,
106∘6′29.7′′ E) is located at Shitouzhai village, about 3 km
east of Ziyun county town in southern Guizhou Province, South China
(Fig. A1). The geologic and paleontologic background of the Shitouzhai
section is described in Appendix A. Its conodont biostratigraphy has been
only partly worked out to date due to sporadic fossil occurrence. Ding and
Huang (1990) identified a few conodont zones that served to demonstrate an
Early to Middle Triassic age for the outcrop. In this study, we detected
three key Early Triassic zonal species in the middle to upper Luolou
Formation: Novispathodus waageni waageni, which ranges from the late
Smithian to early Spathian, and Nv. pingdingshanensis and
Tr. homeri, which are early Spathian in age (Zhao et al., 2007)
(Fig. 1). The first occurrence of Nv. pingdingshanensis is
considered to be a marker of the SSB globally (Zhao et al., 2007) (Fig. 2),
so its appearance in Bed 14 of the study section provides a firm constraint
on the stratigraphic position of the SSB at Shitouzhai. Although the
evolutionary progression of Nv. waageni waageni to
Nv. pingdingshanensis was demonstrated at the better-studied
West Pingdingshan section near Chaohu in Anhui Province (Zhao et al., 2007),
this pattern cannot be established for the present study section owing to the
scarcity of conodont fossils (Fig. 1).
Correlation of the Shitouzhai section in southern Guizhou Province
with the West Pingdingshan section in Chaohu, Anhui Province, South China.
Sources of West Pingdingshan data: conodont zonation (Zhao et al., 2007) and
C-isotope curve (Tong et al., 2007). Correlations between these sections are
based primarily on corresponding features in the C-isotope profiles, although
limited new conodont data for Shitouzhai (central column; n= 4) provide
additional constraints. Fm. and Ht. stand for formation and height,
respectively.
Biostratigraphic and C-isotopic correlations of the
Shitouzhai section with other Smithian–Spathian sections. Note that
Intervals I–IV of δ13Ccarb profiles are recognizable
globally. The standard notation (P2, P3, N3, and N4) for positive and
negative C-isotope excursions of the Early Triassic is according to Song et al. (2013). Data for the Guandao, West Pingdingshan, and Majiashan sections are
from Tong et al. (2007) and from Horacek et al. (2007) for the L'Uomo section. The different color columns represent corresponding conodont zones
from old to young in an ascending order. Question marks represent
problematic conodont biozones in need of further study in the Shitouzhai and
Guandao sections. An., Ind., and Dien. stand for Anisian, Induan, and Dienerian, respectively.
Carbon isotope chemostratigraphy allows the exact placement of the SSB at
Shitouzhai as well as detailed correlation of the study section to
biostratigraphically better-studied sections elsewhere. The δ13Ccarb profile for Shitouzhai shows a pattern of excursions
similar to those of other SSB sections in South China and globally (Fig. 2;
see Song et al., 2013, for a review), indicating that the carbonate carbon
isotope record of the study section was not significantly affected by
diagenesis (Appendix B). The mid- to late Smithian is characterized by a
major negative excursion (N3 of Song et al., 2013), with a minimum δ13C of -3.2 ‰ at Shitouzhai (compared to ca. -1 to
-4 ‰ globally). The SSB is located in the middle of a
rapid positive shift in δ13C having a magnitude ranging from
+3 to +7 ‰ globally. At Shitouzhai, this shift
amounts to +3.5 ‰ and the midpoint of the shift is
located in the upper part of Bed 13, about 50 cm below the base of Bed 14,
thus narrowly constraining the position of the SSB (Fig. 2). There was
limited δ13Ccarb variation during the early Spathian, with
the Shitouzhai study section showing a weak positive drift, whereas most
other sections show a weak negative trend within this interval. All sections
exhibit a large negative δ13Ccarb shift in the mid to late
Spathian, with minimum values ranging from ca. -1 to -4 ‰
(Fig. 2). These δ13Ccarb trends have been well-documented
in Lower Triassic sections from around the world (Payne et al., 2004; Tong
et al., 2007; Horacek et al., 2007; Song et al., 2013; Grasby et al., 2013).
We have correlated the δ13Ccarb profile for Shitouzhai
with that for the biostratigraphically well-constrained West Pingdingshan
section (Tong et al., 2007) (Fig. 1), in which four conodont zones were
recognized within the Olenekian Stage. They are the Nv. w. eowaageni subzone, Nv. w. waageni subzone, Nv.
pingdingshanensis zone, and Triassospathodus homeri zone
(Zhao et al., 2007). The Nv. pingdingshanensis zone is
demarcated by the first occurrences of Nv. pingdingshanensis
and Tr. homeri at its base and top, respectively. At Shitouzhai,
limited fossil occurrences allow the recognition of three of these conodont
zones: the Nv. w. waageni subzone, the Nv.
pingdingshanensis zone, and the Tr. homeri zone (Fig. 1).
The base of the Nv. pingdingshanensis zone (= SSB) also
coincides with a sharp positive δ13Ccarb excursion
that can be correlated globally (Fig. 2).
Methods
Sampling
Large fresh samples, weighing about 3–4 kg each, were collected from outcrop
at the Shitouzhai section. Weathered surfaces and diagenetic veins were
trimmed off, and the remaining sample was crushed into small pieces and
powdered with a rock mill to < 200 mesh for geochemical analysis.
Carbonate carbon isotope analysis
About 80–120 mg of powder was placed in a 10 mL Na glass vial, sealed with
a butyl rubber septum, and reacted with 100 % phosphoric acid at
72 ∘C after flushing with helium. The evolved CO2 gas was
analyzed for δ13C and δ18O using a MAT 253 mass
spectrometer in the State Key Laboratory of Geological Processes and Mineral
Resources at the China University of Geosciences, Wuhan. All isotopic data
are reported as per mille variation (‰) relative to Vienna Pee Dee
belemnite (V-PDB) standard. The analytical precision is better than
±0.1 ‰ for δ13C and ±0.2 ‰ for
δ18O based on duplicate analyses of the national reference standard
GBW-04416 (δ13C = 1.61 ‰).
CAS extraction and sulfur isotope analysis
Carbonate-associated sulfate (CAS) concentrations and isotopes (δ34SCAS) were determined for samples containing
> 30 wt% CaCO3. These samples were powdered, leached of
soluble sulfates in a 10 % NaCl solution, rinsed three times in deionized
water, and dissolved in 3N HCl. The acidified samples were filtered, and an
excess of 1M BaCl2 was added to the filtrate to precipitate BaSO4.
The BaSO4 precipitate was rinsed, filtered, dried, and then combined
with an excess of V2O5 and analyzed for its S-isotope composition
in the State Key Laboratory of Biogeology and Environmental Geology at the
China University of Geosciences, Wuhan. Sulfur isotope compositions are
expressed in standard δ-notation as per mille (‰) variation
with respect to V-CDT (Vienna-Cañon Diablo Troilite), with an analytical
error of ∼ 0.1 ‰ calculated from replicate analyses of samples
and the laboratory standards NBS (National Bureau of Standards) 127
(21.1 ‰), IAEA (International Atomic Energy Agency) SO-5
(0.49 ‰) and IAEA SO-6 (-34.05 ‰). CAS concentrations were
calculated from the mass of recovered BaSO4.
Elemental analysis
The measurement of major and trace element concentrations was carried out in
the State Key Laboratory of Geological Processes and Mineral Resources at the
China University of Geosciences, Wuhan, following the procedure of national
standards (GB/T 14506-2010) and Liu et al. (2008). A Hitachi atomic
absorption spectrophotometer (180-70) and an ultraviolet-visible
spectrophotometer (UV-754) were utilized in major element analysis. An
Aglient 7500a ICP-MS (Inductively coupled plasma mass spectrometer) was used
to analyze trace element concentrations with an average analytical
uncertainty of better than 2 % (RSD, relative standard deviation).
Results were calibrated using the laboratory standards AGV-2, BHVO-2, and
BCR-2. Rare earth element (REE) concentrations were normalized (N) to the
average upper crustal composition of McLennan (2001). In order to calculate
enrichment ratios, lanthanum (La), samarium (Sm), and ytterbium (Yb) were
used as proxies for the light (LREEs), middle (MREEs) and heavy rare earth
elements (HREEs), respectively. The europium anomaly (Eu / Eu*) was
calculated as 2EuN/(SmN+ GdN) and the
cerium anomaly (Ce / Ce*) was calculated as
3CeN/(2LaN+ NdN). The chemical index of
alteration (CIA) was calculated as
Al2O3 / (Al2O3+ K2O + Na2O). This is
a modified form of the original CIA equation (Nesbitt and Young, 1982) that
eliminates CaO from the denominator, which is superior for use in
carbonate-rich sedimentary successions. The Th / Th* ratio, where Th*
represents the average thorium concentration of the upper crust (10.7 ppm;
Bau, 1996), can be used to estimate the fraction of clay minerals in
carbonate units.
Age model and sediment flux calculations
An age model was developed in order to calculate sediment fluxes for the
Shitouzhai section. Age constraints were provided by chemical
abrasion–thermal ionization mass spectrometry (CA-TIMS) studies of U–Pb in
zircons, from which the dates of the Induan–Olenekian and Smithian–Spathian
boundaries were estimated at ∼ 251.25 Ma and ∼ 250.55 Ma, respectively (Ovtcharova et al., 2006), and that of the
Olenekian–Anisian boundary was estimated at ∼ 247.3 Ma (Lehrmann et al.,
2006). These dates yield durations for the Smithian and Spathian substages
of ∼ 0.7 and ∼ 3.25 Myr, respectively. The
age of each sample in the Shitouzhai section was estimated through linear
interpolation between these dated horizons (Fig. 3a). This age model yielded
linear sedimentation rates (LSRs) of 43 and 21 m Myr-1 for
the Smithian and Spathian portions of the study section, respectively.
Sediment bulk accumulation rates (BAR) were calculated as
BAR=LSR×BSD/10,
where an average value of 2.5 g cm-3 was assumed for bulk sediment
density (BSD) and 10 is a coefficient to convert units of m Myr-1 × g cm-3
to g cm-3 kyr-1. The mass accumulation
rates of carbonate and clay (MARcarb and MARclay, respectively) in
the study section were calculated as
MARcarb=BAR×%TIC/12.0,MARclay=BAR×%Al/8.04,
where %TIC and %Al are the concentrations of total inorganic carbon
and aluminum in each study sample, respectively, and the coefficients 12.0
and 8.04 are the concentrations in percent of TIC in pure calcium carbonate
and Al in average upper continental crust, respectively (McLennan, 2001).
(a) Age–depth model and (b) sediment
accumulation rates around the SSB in the Shitouzhai section. BAR,
MARcarb, and MARclay stand for bulk accumulation rate, carbonate
mass accumulation rate, and clay mass accumulation rate, respectively. SSB
stands for Smithian–Spathian boundary.
Discussion
Weathering rate changes
Studies of both modern and ancient carbonates show that a primary seawater
signature is characterized by low ∑REE and
relative HREE enrichment (Webb et al., 2009). However, carbonate sediments
containing even a minor amount of clay minerals tend to acquire a
terrigenous REE signal characterized by high ∑REE and strong LREE or
MREE enrichment (Sholkovitz and Shen, 1995; Bright et al., 2009). At
Shitouzhai, ∑REE exhibits a strong positive correlation with Th (r= +0.97;
Fig. 5a), indicating that REEs came from the detrital clay
fraction not the hydrogenous (seawater) fraction (Zhao et al., 2013).
Moreover, the clay fraction (as estimated from Th / Th*) is substantial,
ranging from ∼ 10 to 30 % of the total sample, which
reflects the argillaceous/muddy character of carbonates in the study
section.
Crossplots of (a) ∑REE vs. Th and
(b) ∑REE vs. Y / Ho. Strong positive covariation
demonstrates derivation of REEs primarily from the terrigenous siliciclastic
(clay-mineral) fraction of the sediment.
All samples at Shitouzhai yield Y / Ho ratios of ∼ 30–35
(Appendix C), which are closer to terrestrial values (∼ 25–30) than to seawater values (44–74) (Bau, 1996; Webb et al., 2009).
∑REE also exhibits a modest negative correlation with Y / Ho (r= -0.65; Fig. 5b).
Thus, a large component of the REEs in the study section is
terrestrially derived, probably through release from clay minerals during
diagenesis. Nearly all Eu / Eu* ratios are in the range of 0.9–1.0 (Appendix C), which are typical of crustal rocks and are consistent with the uptake of
REEs from clay minerals (McLennan, 2001). MREE enrichment is rather strong
(most samples yield SmN / YbN > 1.0; Fig. 4), suggesting
the presence of phosphate in the sediment or the influence of pore waters
previously in contact with phosphate (Kidder and Eddy-Dilek, 1994; Bright et
al., 2009).
All of the detrital proxies from the study section provide evidence of a
major decrease in weathering intensity at the SSB. The age–depth model for
the study section (Fig. 3a) shows that the SSB is characterized by a large
decline in linear sedimentation rates (LSRs) from 43 to 21 m Myr-1 and a proportional decrease in bulk accumulation rates (BAR) from
10.7 to 5.3 g cm-2 kyr-1 (Fig. 3b). The
mass accumulation rates (MAR) of both clays and carbonate also declined
across the SSB, although the decline was larger for clays (∼ 80–90 %)
than for carbonate (∼ 30–40%; Fig. 3b). These
proportional differences reflect the greater concentration of clays in
Smithian beds relative to Spathian beds. The sharp decline in ∑REE
concentrations near the SSB (Fig. 4) is also evidence of a decrease in
clay-mineral content upsection. The CIA has been widely used as a proxy for
chemical weathering intensity in sediment source regions (Nesbitt and Young,
1982; Goldberg and Humayun, 2010). The abrupt decline in CIA values at
Shitouzhai, from ∼ 0.76–0.78 to ∼ 0.70–0.72
(Fig. 4), probably indicates a major decrease in chemical weathering
intensity at the SSB. This interpretation is supported by strong
correlations of CIA with many detrital proxies, including Al (r= +0.87),
ΣREE (r= +0.81), Th / Th* (r= +0.81), and LSR (r= +0.93).
Although changes in CIA potentially can be due to changes in sediment
provenance (e.g., Price and Velbel, 2003), the weak correlation of CIA to
Eu / Eu* (r= -0.21) argues against this interpretation.
All detrital proxies for the Shitouzhai section are thus consistent in
documenting a major decrease in both chemical and physical weathering
intensity at the SSB (Fig. 6). These changes are reflected in lower CIA
values, greatly reduced clay-mineral production, and more limited transport
of siliciclastics to shallow marine systems. Lower bulk sediment fluxes
merely reflect a return to more typical long-term values, however, as the
Griesbachian–Smithian interval of the Early Triassic was characterized by
exceptionally high sediment fluxes and chemical weathering rates (Algeo and
Twitchett, 2010). These weathering-related changes at the SSB are likely to
have been due to a sharp, ∼ 5 ∘C temperature
decrease in the tropics (Sun et al., 2012; Romano et al., 2013). Even the
decline in carbonate flux may have been a consequence of reduced riverine
inputs of Ca2+ and CO32- ions to marine systems, although
other factors such as climatic cooling or changes in oceanic thermohaline
circulation may have influenced marine carbonate production.
Evolution of terrestrial and marine environments during the
late Early Triassic: (a) early Smithian, (b) late Smithian
thermal maximum (LSTM), (c) Smithian–Spathian boundary, and
(d) early Spathian. This model integrates changes in subaerial
weathering rates and oceanic productivity and redox conditions documented in
this study with data regarding paleoclimate variation, terrestrial floral
assemblages, and marine biodiversity patterns from other sources (cited in
text). We infer that the modeled environmental changes were ultimately due
to variation in the eruption rate of the Siberian Traps, although this has not
been proven to date. See text for further discussion.
Oceanic redox variation
The concentrations of redox-sensitive trace elements (e.g., Mo, U, and V)
are low (i.e., close to detrital background values) in all samples from the
study section, although there is a slight increase around the SSB,
especially on a Th-normalized basis (Appendix C). However, there is an even
larger increase in Mn / Th at this level (Fig. 4). Under reducing conditions,
Mn2+ is highly soluble and does not accumulate in substantial amounts
in marine sediments. However, suboxic to oxic conditions commonly result in
Mn enrichment through the accumulation of Mn(II) in carbonates or Mn(III) in
oxyhydroxides (Okita et al., 1988). Strong Mn enrichment is thus common on
the margins of reducing deep water masses (Landing and Bruland, 1987). Mn
enrichment in carbonates is accepted as a good indicator of suboxic
conditions (Rue et al., 1997; Pakhomova et al., 2007). At Shitouzhai, the
Mn / Th profile suggests dominantly anoxic conditions below the SSB (0–18 m)
and suboxic conditions above it (20–37 m), although with a brief interlude
of more reducing conditions during the early Spathian (28–32 m;
Fig. 4).
Cerium (Ce) is the only REE that is affected by oxidation-reduction processes
in the Earth-surface environment. Under reducing conditions, Ce3+ has
the same valence as other REEs and, therefore, is not fractionated relative
to them, yielding Ce / Ce* ratios of ∼ 1.0 (German and Elderfield,
1990). Under oxidizing conditions, Ce4+ is preferentially removed from
solution, yielding local sedimentary deposits with
Ce / Ce* > 1.0, whereas the Ce / Ce* ratio of seawater and
of any hydrogenous deposits incorporating REEs from seawater is
< 1.0 (e.g., 0.3–0.4 in the modern ocean). Thus, Ce is potentially
a good proxy for marine paleoredox conditions, provided that a hydrogenous
signal can be measured (Wright et al., 1987). Terrigenous influence (e.g.,
addition of REEs from clay minerals) will generally cause Ce / Ce* ratios
to converge on 1.0, which is by definition the value for average upper
crustal rocks. In the study section, Ce / Ce* ratios vary from 0.79 to 0.88
(Fig. 4). These moderately high values are nominally indicative of suboxic
conditions. However, the Ce / Ce* ratio was probably heavily influenced by
REEs from the clay fraction of the sediment, making the Ce / Ce* ratio of
any hydrogenous contribution uncertain.
Th / U ratios are useful for paleoredox analysis, owing to the redox-dependent
behavior of U. Under oxidizing conditions, U(VI) tends to form stable
carbonate complexes in seawater (Langmuir, 1978; Algeo and Maynard, 2004).
Under reducing conditions, U(IV) is readily removed to the sediment. Th,
however, is not subject to the influence of redox condition, resulting in
higher Th / U ratios under reducing conditions as aqueous U is lost (Wignall
and Myers, 1988). In the study section, a distinct decrease in the Th / U
ratio at the SSB indicates a shift toward more oxygenated conditions, which
was sustained into the early Spathian (Fig. 4). These results are consistent
with dominantly oxic to suboxic conditions in the study area following the
SSB (Fig. 6).
Significance of C-S isotopic variation at the SSB
Seawater sulfate δ34S rose sharply from ∼ +15 ‰ in the late Permian to > +30 ‰
in the Middle Triassic (Claypool et al., 1980;
Kampschulte and Strauss, 2004), although the pattern of increase during the
Early Triassic has only recently begun to be worked out (Song et al., 2014).
The present study provides the most comprehensive analysis of δ34SCAS variation at the SSB of any study to date. The Shitouzhai
section exhibits a distinct, ∼ 10–15 ‰
negative shift in δ34SCAS that is paired with a
∼ 4 ‰ positive shift in δ13Ccarb (Fig. 4). Both shifts are limited to a narrow interval
around the SSB, probably representing no more than ∼ 75–150 kyr
based on average sedimentation rates for the study section (Fig. 3a).
These two features (i.e., negative covariation and a short event interval)
impose significant constraints on the underlying causes of the isotopic
shifts. Most of the Early Triassic is characterized by positive δ13Ccarb–δ34SCAS covariation, a pattern that is
consistent with control by sediment burial fluxes, i.e., co-burial of
organic carbon and pyrite, linked to variations in marine productivity
and/or redox conditions (Luo et al., 2010; Song et al., 2014). In contrast,
negative δ13Ccarb–δ34SCAS covariation
during a short-term event at the SSB is indicative of oceanographic
controls. Specifically, we hypothesize that cooling-driven reinvigoration
of oceanic overturning circulation led to stronger upwelling, mixing
nutrient- and sulfide-rich deep waters upward into the ocean-surface layer
and causing both enhanced marine productivity (hence higher δ13CDIC) and oxidation of advected H2S (hence lower δ34Ssulfate) (Fig. 6). Such an oceanographic process was
inherently transient, lasting only until the nutrients and sulfide that had
accumulated in the deep ocean during the Griesbachian–Smithian interval of
intense oceanic stratification (Song et al., 2013) became depleted. The same
process was inferred for the latest Spathian by Song et al. (2014), an
interval also characterized by short-term negative δ13Ccarb–δ34SCAS covariation (Song et al., 2014,
their figure 6) and linked to global climatic cooling (Sun et al., 2012).
These considerations underscore the fundamental significance of the SSB,
which represents the termination of the Early Triassic hyper-greenhouse
climate and the reinvigoration of global-ocean overturning circulation (Fig. 6).
Causes and consequences of the SSB event
Oceanographic changes at the SSB had a major effect on contemporaneous
marine biotas. Several invertebrate clades, including ammonoids, conodonts,
and foraminifera, appear to have suffered severe losses of biodiversity at
this time (Orchard, 2007; Stanley, 2009; Song et al., 2011). Ammonoids
diversified greatly during the Griesbachian to Smithian but underwent a
major evolutionary turnover at the SSB, followed by a stepwise increase in
biodiversity in the early to middle Spathian (Brayard et al., 2009).
Conodonts show a similar pattern, with a rapid radiation in the
early to middle Smithian terminated by a severe extinction at the SSB,
followed by a second radiation in the early to middle Spathian (Orchard,
2007). Changes in biodiversity were mirrored by changes in body size. Chen
et al. (2013) documented a brief but significant size reduction among
conodonts, coinciding with the late Smithian thermal maximum (Sun et al.,
2012), based on bulk sample analysis from an outcrop section in Guizhou
Province, southwestern China. Conodonts remained diminutive during the SSB
transition and the earliest Spathian and then underwent a stepwise size
increase during the early to middle Spathian (Chen et al., 2013).
Although literature surveys show that marine clades such as conodonts,
ammonoids, and foraminifera experienced a sharp decline in diversity at the
SSB (Orchard, 2007; Stanley, 2009; Song et al., 2011), this pattern may be
biased by data binning effects. In fact, an examination of the stratigraphic
distribution of these marine clades in actual geological sections suggests
that diversity losses occurred slightly prior to the SSB (Zhao et al., 2007;
Song et al., 2011; Zakharov and Popov, 2014) and were probably associated
with the late Smithian thermal maximum (Sun et al., 2012; Romano et al.,
2013; Fig. 6) rather than the Smithian–Spathian boundary itself. The
affected marine clades also did not recover immediately when climatic and
environmental conditions ameliorated abruptly at the SSB but, rather,
underwent a stepwise recovery during the early to middle Spathian (Orchard,
2007; Stanley, 2009; Brayard et al., 2009).
The SSB was characterized by a major change in terrestrial flora.
Lycopsid-dominated assemblages were replaced by conifer-dominated or mixed
lycopsid-conifer vegetation, as indicated by palynological data from Pakistan
(Hermann et al., 2011), Norway (Galfetti et al., 2007; Hochuli and Vigran,
2010), and central Europe (Kurscher and Herngreen, 2010). A similar floral
change was reported from the Spathian–Anisian boundary in Hungary (Looy et
al., 1999), suggesting some variation in the timing of terrestrial floral
recovery in different regions of the world. Macrofloral fossil evidence
indicates a more volatile record of vegetation change, with multiple
short-term expansions of lycopsids from tropical regions temporarily
displacing conifers during the Olenekian (Retallack et al., 2011; Hochuli et
al., 2010; Looy et al., 2001). These inferences are supported by biomarker
and biogeochemical studies. Saito et al. (2013) reported that sediments of
Griesbachian to Smithian age yield carbon/nitrogen (C / N) ratios
< 10 and contain abundant retene, simonellite, and dehydroabietan,
which are interpreted to have been sourced from lycopsids and/or bryophytes.
After the SSB, sediments yield C / N ratios > 10 and exhibit a
large increase in pimanthrene abundance, suggesting the dominance of
terrestrial floras by conifers. As a result, a highly diverse coniferous
flora became widely reestablished around the SSB, replacing the lycopsid- and
fern-dominated disaster-type vegetation that had dominated the Griesbachian
to Smithian interval (Saito et al., 2013; Fig. 6).
The SSB was also characterized by major environmental changes. Strong
climatic cooling has been inferred from both faunal (Galfetti et al., 2007)
and oxygen-isotope evidence (Sun et al., 2012; Romano et al., 2013). Changes
in oceanic circulation appear to have occurred at the same time. Saito et
al. (2013) interpreted an increase in extended tricyclic terpane ratios (ETR)
around the SSB as being due to a shift from limited to vigorous overturning
circulation (Fig. 6). These climatic and oceanographic changes were probably
linked: an increase in the intensity of global meridional circulation would
have been a natural consequence of climatic cooling (e.g., Rind, 1998),
leading to more vigorous deepwater formation in high-latitude regions (Kiehl
and Shields, 2005).
The environmental and climatic changes documented at Shitouzhai reinforce
observations made in other SSB sections globally and, thus, serve to
demonstrate that these changes were widespread and characteristic of the
SSB. We propose that all of the changes in our model (Fig. 6) were due to a
cooling event that commenced following the LSTM and that continued strongly
across the SSB. In particular, we infer that cooling led to the reinvigoration
of global-ocean overturning circulation. It should be noted that we are not
envisioning complete ocean stagnation during the preceding
Griesbachian–Smithian interval, which is unlikely, based on physical
oceanographic principles (e.g., Kiehl and Shields, 2005), but, rather, a
strong slowing of overturning circulation that led to a buildup of nutrients
in the deep ocean (Fig. 6). The reinvigoration of global-ocean circulation at
the SSB flushed this buildup of nutrients back into the ocean-surface layer,
triggering a transient increase in marine productivity and an expansion of
thermoclinal anoxia that lasted until this deepwater nutrient source was
depleted. The brevity of the SSB anoxic event at Shitouzhai, which lasted
∼ 75–150 kyr, is consistent with such a mechanism. This
mechanism also accounts for the abrupt, large positive shift in δ13Ccarb at the SSB, which was due to a productivity-related
increase in organic carbon burial rates (Fig. 6).
The ultimate cause of the SSB event is uncertain. Given that the onset of
the Permian–Triassic boundary crisis has been firmly linked to the initiation of
the main eruptive phase of the Siberian Traps large igneous province (STLIP)
(Renne et al., 1995; Kamo et al., 2003) and that the Early Triassic was an
interval of repeated environmental disturbances (Algeo et al., 2011;
Retallack et al., 2011) and elevated global temperatures (Sun et al., 2012;
Romano et al., 2013) linked to volcanogenic greenhouse gas emissions
(Retallack and Jahren, 2008; Black et al., 2012), the obvious explanation
for the SSB is a reduction in the intensity of magmatic activity in the
STLIP source region (Fig. 6). The available radiometric age data for the
Siberian Traps, although sparse, are consistent with this possibility. U–Pb
dating of perovskites in the early Arydzhangsky flow and zircons from the
late Delkansky silicic tuff of extrusive suites in the Maymecha–Kotuy region
suggests that the STLIP flood basalt eruptions commenced at 251.7 ± 0.4 Ma
and ended at 251.1 ± 0.3 Ma, i.e., representing an interval of ∼ 600 kyr
(Renne et al., 1995; Kamo et al., 2003). However, an Ar–Ar date of
250.3 ± 1.1 Ma was obtained for the final stage of extrusive volcanism
at Norilsk, the core area of the STLIP (Reichow et al., 2009; see also the review of evidence for a late eruptive stage by Ovtcharova et al., 2006).
The more critical issue, in any case, is the duration not of flood basalt
eruptions but of intrusive magmatism in the West Siberian Coal Basin, which
was probably the main source of volcanogenic greenhouse gases (Retallack and
Jahren, 2008; Black et al., 2012). Reichow et al. (2009) reported ages for
STLIP-related intrusives spanning several million years, which is consistent
with the hypotheses that large-scale intrusive activity continued at least
until the SSB and that cessation of most such activity at the SSB was
responsible for contemporaneous climatic cooling (Sun et al., 2012; Romano
et al., 2013). Further work on the chronology of the STLIP will be needed to
conclusively evaluate controls on the SSB event.