Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
Contents lists available at ScienceDirect
Palaeogeography, Palaeoclimatology, Palaeoecology
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / p a l a e o
Mantle plume: The invisible serial killer — Application to the Permian–Triassic
boundary mass extinction
Ezat Heydari a,⁎, Nasser Arzani b, Jamshid Hassanzadeh c
a
b
c
Department of Physics, Atmospheric Sciences, and Geoscience, Jackson State University, P.O. Box 17660, Jackson, MS 39217, USA
Geology Department, University of Payame Nour, Kohandej Street, Esfahan, Iran
School of Geology, College of Sciences, University of Tehran, Tehran, Iran
A R T I C L E
I N F O
Article history:
Received 3 July 2007
Received in revised form 28 March 2008
Accepted 4 April 2008
Keywords:
Permian
Triassic
Mantle Plume
Iran
Shahreza
Mass extinction
A B S T R A C T
The Earth experienced a severe mass extinction at the Permian–Triassic boundary (PTB) about 252 million
years ago. This biological catastrophe was accompanied by major changes in geochemical composition of the
atmosphere and ocean and the appearance of sedimentary features which had not occurred since the
Precambrian time. The eruption of the largest continental flood basalt, the Siberian Traps, overlapped this mass
killing. Many hypotheses have been proposed but no definitive conclusion currently exits. Here we present
characteristics of three sections from Iran and China and propose that an active mantle plume initiated a
series of processes which led to the mass mortality and produced major sedimentological, mineralogical, and
geochemical changes observed in the transition from the Paleozoic to the Mesozoic.
The injection of mantle plume-related igneous dike swarms into the continental margin facilitated the release
of massive amounts of CH4 primarily from the dissociation of marine gas hydrates and secondarily from the
maturation of organic-rich sediments and fracturing of hydrocarbon reservoirs. The bulk of the CH4 was
aerobically oxidized in the water column producing dissolved CO2 with low δ13C values. This CO2-saturated
seawater became acidic to the point of dissolution of shelf carbonates promoting precipitation of siliciclasticrich strata in the transition from the Permian to the Triassic. Methane-derived CO2 also lowered carbon isotopic
composition of seawater leading to the observed decline in δ13C composition of organic and inorganic marine
carbon at the PTB.
Gas-charged oceans released large volumes of CO2 and CH4 into the atmosphere which created a severe global
warming (the end-Permian inferno) causing the release of additional CH4 from the dissociation of polar gas
hydrates. These events lowered δ13C compositions of terrestrial carbon. Simultaneously, feeder dikes from the
mantle plume formed the Siberian Traps flood basalt.
Marine mass extinction was the result of a change in seawater composition due to the injection and oxidation of
CH4 in the water column causing low pH, high concentrations of CO2, Ca2+ and HCO3−, and low CO32− values.
Combined with a hot seawater temperature, these changes made calcification of marine organisms difficult and
produced major physiological crisis including reduced metabolic rates, high sensitivity to environmental stress,
and hampered growth and reproduction. Terrestrial mass extinction can be attributed to severe global
warming and soil acidification produced by increased atmospheric CO2, acid rain that was generated by SO2
derived from the Siberian trap eruption, and loss of habitat.
Cessation of the plume activity during Early Triassic stopped the release of CH4 into the ocean and terminated
continental flood basalt eruption ending the environment of death on land and in sea. The cut off of CO2
production in the ocean instantly increased carbonate saturation of seawater resulting in extensive seafloor
cementation. It also resulted in the deposition of marine carbonates by microbial activities in the hostile postextinction environment. From the trigger to recovery, the perturbation which included the end-Permian mass
mortality could have lasted for at least 2 Myr.
Several major mass extinctions of the Phanerozoic are temporally accompanied by flood basalt eruptions. So
far, these two events have been interpreted in a cause-and-effect relation: flood basalt eruption causes mass
extinction. This study proposes that flood basalts and their time correlative biological crises are themselves the
consequence of a complex perturbation caused by mantle plume activities. If so, major disturbances in the near
surface of the Earth are ultimately controlled by changes in the mantle.
⁎ Corresponding author.
E-mail addresses: ezat.heydari@jsums.edu (E. Heydari), arzan2@yahoo.com (N. Arzani), jamshidh@khayam.ut.ac.ir (J. Hassanzadeh).
0031-0182/$ – see front matter © 2008 Elsevier B.V. All rights reserved.
doi:10.1016/j.palaeo.2008.04.013
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E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
The physical sedimentological observations presented here combined with previous paleontological evidence
cast doubt on other interpretations which use geochemical variables, numerical modeling, biomarkers, and Ce
anomaly to suggest that Late Permian ocean was anoxic all the way to the photic zone.
© 2008 Elsevier B.V. All rights reserved.
1. Introduction
3. Sedimentology and geochemistry of PTB strata
A series of extraordinary perturbations affected the Earth during
the Late Permian to Early the Triassic time interval (257–247 Myr
ago) (Payne et al., 2004). These disturbances began with the mass
extinction and the eruption of the Emeishan continental flood basalt
at the end of the Guadalupian (257 Myr ago) (Stanley and Yang, 1994;
Zhou et al., 2002; Isozaki et al., 2007a). This was followed by the
most devastating biological crisis of the Phanerozoic and the
formation of the Siberian Traps flood basalt at the Permian–Triassic
boundary (PTB) about 252 Myr ago (Renne and Basu, 1991; Erwin,
1993; Retallack, 1995, 1999; Twitchett et al., 2001; Reichow et al.,
2002; Mundil et al., 2004; Payne et al., 2004; Erwin, 2006). Changes
continued during the Early Triassic as shown by several major negative and positive shifts in δ13C composition of marine carbonates
(Payne et al., 2004; Richoz, 2006; Galfetti et al., 2007; Horacek et al.,
2007a,b).
Of these perturbations, the one at the PTB has received the most
attention because of its consequence for life on Earth (Retallack, 1995,
1999; Twitchett et al., 2001; Payne et al., 2004; Erwin, 2006). Bolide
impact, supernova explosion, changes in seawater salinity, marine
anoxia, volcanism, oceanic overturn, hydrogen sulfide expulsion,
severe climate change, dissociation of gas hydrates, and coal
metamorphism are scenarios proposed to explain the events that
occurred at the PTB (Fischer, 1965; Xu et al., 1985; Stanley, 1988;
Wignall and Hallam, 1992; Kajiwara et al., 1994; Ellis and Schramm,
1995; Renne et al., 1995; Knoll et al., 1996; Bowring et al., 1998;
Woods et al., 1999; Becker et al., 2001; Kaiho et al., 2001; Heydari
et al., 2003; Heydari and Hassanzadeh, 2003; Kump et al., 2005;
Erwin, 2006; Retallack et al., 2006). But no definitive conclusion exits
as all hypotheses have been debated (Erwin, 1993; Erwin et al., 2002;
Koeberl et al., 2004; Erwin, 2006). This is clearly demonstrated by the
diversity of interpretations expressed in the most recent publications
dedicated to this topic (Baud et al., 2007; Farabegoli et al., 2007; Haas
et al., 2007; Isozaki et al., 2007b).
The current state of knowledge and the lack of a unified model
necessitate additional studies of strata encompassing the transition
from the Paleozoic to the Mesozoic. The goals of this investigation are
as follow: (1) to describe features of three PTB sections in Iran and
China (Fig. 1), (2) to concentrate on detailed lithological, petrographical, and geochemical characteristics of the Shahreza section of Iran,
and (3) to elaborate on the proposed hypothesis that a mantle plume
was responsible for the PTB mass extinction and other events that
took place at that time.
3.1. Chaotian section, China
In the Chaotian section of China, the uppermost Permian is
represented by the Dalong Formation which is subdivided into four
units: “A”, “B”, “C”, and “D” (Isozaki et al., 2004, 2007b) (Fig. 1B). All four
units are grey to black in color and thin-bedded (5–15 cm). Units “A”
(N0.9 m) consists of calcareous mudstone. Unit “B” (2.8 m) is lime
mudstone. Units “C” (1.4 m) and “D” (2.3 m) are wackestone (Fig. 1B). All
four units are fossiliferous and contain bivalves, gastropods, brachiopods, ammonoids, ostracods, radiolarians, and conodonts (Isozaki et al.,
2007b). Most importantly, bioturbation is absent in units “A” and “B”,
appears in unit “C”, but becomes abundant in unit “D” (Isozaki et al.,
2007b). That is, bioturbation increases upward toward the end-Permian
mass extinction that occurred at the top of the unit “D” (see below). Total
organic carbon content (TOC) values are about 1–4 wt.% in unit “A” and
decrease to 0.02 to 0.14 wt.% in unit “D”. In other words, TOC decreases as bioturbation increases.
The top of the Dalong Formation represents the Event Horizon
which is the stratigraphic position of the end-Permian mass mortality
(Fig. 1B). The overlying Feixianguan Formation displays three units:
“E”, “F”, and “G” (Isozaki et al., 2007b). The basal “E” unit (1.4 m)
consists of olive-gray, faintly laminated marl. It is barren of fossil
except for small-size bivalves and ammonite (Fig. 1B). Carbonate
content of unit “E” is about 41%. The layer is also rich in total iron
relative to the rest of PTB interval. Unit “E” has very low TOC values:
0.06 to 0.6 wt.% (Isozaki et al., 2007b).
The top of the unit “E” marks the PTB defined by the first appearance
of H. parvus conodont (Fig. 1B). The overlying unit “F” (1.7 m) consists of
faintly laminated grey, lime mudstone. Its carbonate content ranges
from 77 wt.% to 95 wt.%. TOC remain low in unit “F” with values ranging
from 0.02 wt.% to 0.19 wt.%. The overlying unit “G” (N1.4 m) consists of
dark grey lime mudstone with faint lamination. TOC content reaches its
lowest values in unit “G”: 0.04 to 0.07 wt.% (Isozaki et al., 2007b).
Unit “E” was originally called the “boundary clay” by Isozaki et al.
(2004). We suggest that the term “boundary clay” should be avoided
when studying the PTB strata, because these strata are neither clay nor
they occur at the boundary (Fig. 1B). Instead, we use the term
Transition Zone (TZ) to distinguish strata that were deposited after the
Event Horizon and before the first appearance of H. parvus which
marks the beginning the Triassic Period (Fig. 1B). The characteristics of
Transition Zone strata are of particular interest because they formed
immediately after the end-Permian mass mortality.
2. Methods
3.2. Julfa section, Iran
Sedimentological characteristics of each layer were studied
through common field methods. Petrographic observations were
made on polished and stained thin sections. Crystal sizes were
determined on enlarged transmitted-light photographs. An unaltered
portion of samples was cut with a trim saw, washed thoroughly, dried,
powdered, and analyzed for bulk mineralogy by X-ray diffraction. For
Sr concentration, 100 mg of the powder was dissolved in 0.5 N HCl and
the solution was analyzed by inductively coupled plasma atomic
emission spectroscopy (ICP-AES) with a resolution of ±1 ppm. For δ13C
compositions, 10 mg of sample was reacted with orthophosphoric acid
and the resulting CO2 was analyzed. The precision of analysis is ±0.2‰
PDB (Peedee Belemnite).
The first comprehensive study of the Julfa section was conducted
by Stepanov et al. (1969) who subdivided the Permian and the Triassic
strata into seven units: “A” to “G”. Teichert et al. (1973) presented one
of the most comprehensive paleontological investigations of this
locality. Baud et al. (1989) provided carbon isotopic composition and
Kakuwa and Matsumoto (2006) analyzed Ce anomaly and δ13C values
of this section. The lithostratigraphic and paleontologic descriptions
presented here and shown in Fig. 1C are from Kozur (2003) and
Szurlies and Kozur (2004), and Kozur (2007).
The uppermost Permian Ali Bashi Formation consists of more than
5 m of red to brownish, fossiliferous, highly bioturbated to nodular,
micritic limestone to marl with abundant ammonoids, brachiopods, and
E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
Fig. 1. (A). Location of the three sections discussed in this study are shown on the Late Permian paleogeographic reconstruction (Golonka, 2000). (B). Stratigraphic column of the Chaotian section in China (modified from Isozaki et al., 2004,
2007b). Biozones include: a = Clarkina subcarinata, b = Clarkina deflecta, c = Clarkina carinata, d = Pseudogastrioceras spp., e = Pentagoceras spp., and f = Pleuronodoceras spp. EH = Event Horizon. PTB = Permian–Triassic boundary. (C). Stratigraphic
column of the Julfa section in Iran (modified from Szurlies and Kozur, 2004 and Kozur 2007). Carbon isotope data from Kakuwa and Matsumoto (2006). Biozones include: a = Clarkina iranica, b = C. yini, C. zhangi, C. changxingsis, C. deflecta,
c = Neogondolella carinata, d = Pseudogastrioceras sp. EH = Event Horizon. PTB = Permian–Triassic boundary. (D). Stratigraphic column of the Shahreza section in Iran (this study). Conodont biozones (from Korte et al., 2004a) include: a = Clarkina
iranica, b = C. yini, C. zhangi, C. changxingsis, C. deflecta, c = Clarkina nodbsa, and d = Clarkina bachmani. See the text for the lithological units. EH = Event Horizon. PTB = Permian–Triassic boundary.
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Fig. 2. Graphs show the stratigraphy (A), variations of non-carbonate fraction (B), changes in calcite fraction (C), Sr concentration (D), and δ13C signature of the uppermost Permian and the lowermost Triassic strata of the Shahreza section. Data
in B and C are from X-ray diffraction analyses. (A). The uppermost Permian to the lowermost Triassic stratigraphic column of the Shahreza section of Iran (see Fig. 1 for explanations). EH = Event Horizon. PTB = Permian–Triassic boundary.
TZ = Transition Zone. (B). Non-carbonate component is very low in the uppermost Permian strata and increases rapidly in the Transition Zone (TZ) interval and reaches its lowest values in the lowermost Triassic layers. (C). The percentage of
calcite is very high in the uppermost Permian but declines in the Transition Zone (TZ) interval. The lowermost Triassic is a nearly pure limestone. (D). Sr concentration remains low and invariant in the uppermost Permian interval (average
575 ppm) and in the Transition Zone (average 480 ppm), but it increases (average 750 ppm) and fluctuates wildly in the lowermost Triassic strata. (E). The δ13C composition has high values (+4‰ PDB) in the uppermost Permian Hambast
Formation. However, δ13C value begins to decline at about 9 m below the top of the Hambast Formation reaching +1‰ PDB at the Event Horizon (EH). The decline continues in the Transition Zone (TZ) interval reaching a value of − 0.5‰ PDB
near the PTB. The lowest value of −0.7‰ PDB is reached at about 4 m above the PTB. The δ13C composition increases to a value of +1‰ PDB at the top of the measured section.
E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
few deep-water corals (Fig. 1C). The top of the Ali Bashi Formation marks
the Event Horizon (Fig. 1C).
The lowermost part of the overlying Elika Formation begins with 0.9 m
of red, bioturbated shale to silty mudstone which is devoid of calcareous
fossils. This layer was called the “boundary clay” by Kozur (2003) and
Szurlies and Kozur (2004). The red shale is overlain by 0.15 m of pink to
red argillaceous lime mudstone followed by 0.3 m of red to brown to
green siltstone and marl, the top of which coincides with the first
appearance of the H. parvus conodont (Kozur, 2003; Szurlies and Kozur,
2004; Kozur 2007). As in the Chaotian section, the stratigraphic interval
between the top of the Ali Bashi Formation and the appearance of the
H. parvus is designated as the Transition Zone (TZ) in this study (Fig. 1C).
The lowermost Triassic begins with 0.8 m of yellow to pink to blue
grey limestone and marl (Fig. 1C). This is followed by N 5 m violet
weathered, yellowish to brownish, grey to pink laminated crinoidbearing limestone (Fig. 1C).
The δ13C compositions of the uppermost Permian are about +3‰
PDB (Kakuwa and Matsumoto, 2006). This value begins to decline at
about 3 m below the Event Horizon reaching −0.5‰ PDB there (Fig. 1C).
After a 0.8‰ PDB increase, the δ13C reaches its lowest composition
of −1‰ PDB about 1 m above the Permian–Triassic paleotological
boundary (Fig. 1C). Kakuwa and Matsumoto (2006) also provide Ce
anomaly of this section which will be discussed below.
3.3. Shahreza section, Iran
The overall stratigraphy and isotopic composition of the Shahreza
section was originally presented by Heydari et al. (2001). Subsequent
151
investigations by Korte et al. (2004a) and Kozur (2007) provided
conodont zonation and δ13C composition of this locality. The present
study provides the most detailed sedimentological and geochemical
analysis of the PTB strata of the Shahreza section.
The uppermost Permian Hambast Formation (bed 1) consists of
red, fossiliferous, highly bioturbated to nodular, micritic limestone,
consisting of over 95% calcite (Figs. 1D, 2B, C). Micrite crystals are
small (5–10 μm) and show no signs of recrystallization (Fig. 3A). Sr
concentrations are uniformly low, averaging 575 ppm (Fig. 2D).
The δ13C compositions are high (4‰ PDB) at the base but begin a
nearly 5‰ PDB decrease at about 9 m below the top of the Hambast
Formation (Fig. 2E).
The top of the Hambast Formation represents the Event Horizon
(Fig. 1D). The overlying informally-named Shahreza Formation begins
with bioturbated red shale and marl (Bed 2, 0.7 m) which are barren of
fossils and poor in carbonate (Fig. 2B, C). It is overlain by bed 3 (0.6 m)
which is composed red to brownish paper-thin marl to argillaceous
limestone (Fig. 2B, C). Bed 4 (0.7 m) consists of red shale and marl
(Fig. 1D). The H. parvus conodont marking the PTB appears at about
2 m above the Event Horizon (Fig. 1D). Therefore, Beds 2–4 are
considered to be the Transitional Zone (TZ) layers (Fig. 1D). Sr values
remain low in the TZ interval, averaging 480 ppm (Fig. 2D). The overall
decline in δ13C compositions which had begun in the underlying
Hambast Formation continues in the TZ interval, reaching 0‰ PDB at
the paleontologically defined PTB (Fig. 2E) (Fig. 1D).
The lowermost Triassic strata consist of tan-colored limestone with
less than 5% impurities (Fig. 2B, C). Their deposition began with a layer
that contains fan-shaped calcite crystals each 10–20 cm long (bed 5),
Fig. 3. (A). Photomicrograph (plane polarized light) of the uppermost Permian Hambast Formation of the Shahreza section (bed 1). The limestone consists of small crystals (5–10 μm), lacks
recrystallization fabrics, and has low Sr concentrations (575 ppm, see Fig. 2E). These criteria suggest that the precursor mineralogy of the uppermost Permian limestone was calcite. (B). A
typical photomicrograph (plane polarized light) of the lowermost Triassic strata (beds 5–9) Shahreza Formation. Cloudy-looking areas were grains whose original fabric was totally
obliterated and replaced by large calcite crystals (10–150 μm). These limestones also show high and large fluctuation in Sr concentration which average 775 ppm (see Fig. 2D). These criteria
suggest that the original mineralogy of the lowermost Triassic limestone was aragonite. (C). Photograph shows meter-high mounds in the lowermost Triassic strata. Field and petrographic
observations indicate that the mounds were constructed by large (10–50 cm) radial calcite crystals. The overlying strata conform to the shape of the mound. These features indicate that the
mound formed a synsedimentary growth on the seafloor. Such seafloor growth was typical of the Precambrian Earth (see Pruss et al., 2006 for review). (D). Field sketch shows internal
structure of the mound next to person in photograph “C”. Note that the mound is constructed by crystal fans some of which extend across the entire thickness of the mound.
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followed by alternating marl and limestone (bed 6), and thin- to
medium bedded limestone (bed 7) (Fig. 1D). These limestones are
dominated by structure-less grains 0.2–1 mm in size (possibly
peloids) and some fossil fragments both of which are totally
recrystallized beyond recognition (Fig. 3B). Small (1–5 cm) stromatolitic domes are common. Laminations are faint, irregular, noncontinuous and very similar to those formed through binding of
sediments by microbial activities.
Most importantly, the lowermost Triassic limestone contains
meter-sized mounds constructed by 10–50 cm-long radial fans of
calcite crystals (Fig. 3C, D). Succeeding layers conform to the shape of
these mounds suggesting that mounds formed by synsedimentary
precipitation that grew on the seafloor (Fig. 3C).
In contrast to the uppermost Permian strata, the lowermost
Triassic limestone consists of large crystals (10–150 μm) and shows
total obliteration of the original fabric (Fig. 3B), both indicative of postdepositional recrystallization. Sr concentrations are high and show
large fluctuations between 400 ppm and 1800 ppm, with an average of
775 ppm (Fig. 2D). After a sharp increase just above the PTB, the δ13C
compositions continue its decline reaching a minimum of −0.7‰ PDB
about 4 m above the PTB, and then increase reaching a value of +1‰
PDB at the top of the studied section (Fig. 2E).
4. Discussions
4.1. Depositional environments
Paleontological investigations suggest that the Shahreza, Julfa, and
Chaotian sections are complete with no hiatus present (Kozur, 2003;
Szurlies and Kozur, 2004; Korte et al., 2004a; Isozaki et al., 2007b;
Kozur, 2007). Therefore, the sedimentary layers of these three sections
have recorded geological process that occurred in the open ocean
during the end-Permian crisis.
4.1.1. The Uppermost Permian: before and during the great dying
The uppermost Permian (Changhsingian) strata representing the
interval prior to the Event Horizon in the Chaotian, Julfa, and Shahreza
sections show major similarities. The lithofacies of this interval is
characterized by micritic limestone deposited in deep-water: 100–
200 m in the Shahreza and Julfa sections and in the slope environment
in the Chaotian locality (Isozaki et al., 2007b; Kozur, 2007). A similar
conclusion was reached by Heydari et al. (2003) and Korte et al.
(2004b) for the uppermost Permian of the Abadeh section.
Most importantly, the Changhsingian limestones of the Julfa and
Shahreza sections are highly bioturbated to the point of becoming
nodular before and during the time when mass killing was taking
place. Units “A” and “B” of the Chaotian section are laminated, but
bioturbation appears in unit “C” and becomes abundant in unit “D”. In
addition, fossil abundance increases toward unit “D”; whereas TOC
content decreases in the same interval. These observations suggest
that open marine remained oxygenated before and during the mass
killing all the way to the slope environment as was also concluded
by Korte et al. (2004b) and Kozur (2007) from the study of other
sections.
4.1.2. The uppermost Permian: the Transition Zone
“Normal” marine processes represented by the uppermost Permian apparently came to an end after the Event Horizon. The
Transition Zone (TZ) strata are characterized by siliciclastic shale
and marl in all three sections (Fig. 1). The difference is their color: red
in the Julfa and Shahreza sections, gray in the Chaotion locality. More
importantly, carbonate-poor red shales (marine red beds) that
immediately overly the Event Horizon in the Julfa and the Shahreza
sections (Bed 2 in Fig. 1D) are non-laminated meaning that the
oxygenation continued even after the mass extinction had occurred
(Kozur, 2003; Szurlies and Kozur, 2004; Kozur, 2007). The TZ interval
is faintly laminated in the Chaotian section (Isozaki et al., 2007b)
(Fig. 1B). But bioturbation and laminations seem to alternate in the
upper portion of TZ strata in the Julfa and the Shahreza sections (Fig. 1D).
The faint lamination of the TZ strata in these three sections cannot be
attributed to anoxia, because of the very low TOC content of these layers
and the red color in the Julfa and Shahreza sections. Very likely, the
lamination is preserved due to the absence of burrowing organisms.
4.1.3. Early Triassic: a new world
The Early Triassic interval in the three sections is slightly different.
It consists of gray, faintly laminated lime mudstone in Chaotian
section; it is composed of pink to red, skeletal limestone in the Julfa
locality; and it is made up of white limestone dominated by microbial
features and mounds consisting of crystal fans in the Shahreza area.
Variations in the Lower Triassic lithofacies in these three sections may
be due to differences in water depth and local environmental process.
However, investigations show that sedimentologic characteristics
observed at the Shahreza section was typical of the Early Triassic
time. This is because features similar to the Shahreza section are
reported from Turkey, Armenia, southwestern United States, Hungry,
Italy, south China, and Greenland (Sano and Kakashima, 1997; Woods
et al., 1999; Kershaw et al., 1999; Lehrmann, 1999; Lehrmann et al.,
2003; Pruss and Bottjer, 2004; Baud et al., 2005; Pruss et al., 2005,
2006; Lehrmann et al., 2007; Baud et al., 2007; Woods et al., 2007).
Excellent review of the Lower Triassic sedimentation was presented by
Pruss et al. (2006) and Baud et al. (2007).
Calcified microbial framework, similar to those found in the
Chinese sections (Lehrmann, 1999; Lehrmann et al., 2003, 2007),
were not seen in the Lower Triassic strata of the Shahreza section.
Small stromatolitic features are the clearest evidence of microbial
deposition in the Shahreza section. However, the origin of structureless, highly recrystallized, micritic lumps (possibly peloids) is
uncertain. With all likelihood, these peloids were also microbially
produced. The meter-size mounds with large crystal fans were
produced by direct precipitation of carbonates on the seafloor, a type
of carbonate deposition that was dominant during the Pre-Cambrian
(see Pruss et al., 2006).
4.2. Implications of sedimentological observations
Perhaps the most important implication of the sedimentological
observations of the three sections is the lack of definitive feature
indicative of anoxia before and during the end-Permian mass
mortality. This includes the absence of well-defined varved-type
lamination and the associated organic-rich sediments. Therefore, the
mass killing occurred while the water column was oxygenated all the
way to the middle of the slope area. This interpretation is also
supported by detailed paleontological investigations indicating that
deep-water strata of the uppermost Permian contain abundant
organisms many of which survived the mass extinction and thrived
in deep-water settings of the earliest Triassic (Twitchett et al., 2004;
Chen et al., 2006a,b; Kozur, 2007).
Our conclusion regarding the state of water column oxygenation at
the time of the end-Permian mass mortality is in direct conflict to
many investigations which imply water column anoxic during this
time interval. The most highly cited of the oceanic anoxia scenarios is
that of Knoll et al. (1996) who adapted an oceanographic model
developed by Hoffman et al. (1991) to explain marine geochemical
variations of the Phanerozoic ocean including those at the PTB. The
model proposes that the Permian ocean consisted of a stratified water
column consisting of a bottom anoxic water mass and the overlying
oxygenated water mass. Knoll et al. (1996) used non-exiting polar icecap to induce mixing of these two water columns and initiate the mass
extinction.
The “laminated” black shale on which Wignall and Hallam (1992)
based their anoxia-driven kill model was deposited in brackish coastal
E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
lagoon and not in an open ocean system (see Heydari, 2005). The dark
color chert which Isozaki (1997) used to propose his super-anoxia
interpretation of the PTB interval could have been deposited in a backarc basin not in deep open ocean (Zhang et al., 2001). Additionally,
sedimentological results of Chaotian section discussed by Isozaki et al.
(2007b) indicate oxygenation actually increased toward the endPermian mass extinction in the slope environment contradicting the
frequently-cited “super-anoxia” model of Isozaki (1997).
Furthermore, the physical sedimentological observations presented here combined with previous paleontological evidence cast
doubt on other interpretations which use geochemistry (Riccardi
et al., 2006), numerical modeling (Kump et al., 2005), biomarkers
(Grice et al., 2005; Hays et al., 2007), and Ce anomaly (Kakuwa and
Matsumoto, 2006) to suggest that the end-Permian ocean water
column was anoxic all the way to the photic zone.
The Transition Zone (TZ) layers are poor in carbonate and consist of
siliciclastic shale and marl in all three sections (Fig. 1), suggesting a
reduction in carbonate production worldwide the cause of which will
be discussed in the following paragraphs. The TZ strata that
immediately overly the Event Horizon in the Shahreza and the Julfa
sections are bioturbated indicating that oxygenation continued even
after the mass killing had ended in these two areas. Faint laminations
occur in TZ and Lower Triassic strata (Fig. 1). We do not consider these
laminations to be indicative of anoxic condition because they contain
very low TOC content and are red in color in the Shahreza and Julfa
sections. One would expect high TOC values if the Late Permian Ocean
were anoxic all the way to the photic zone and were infested with
bacteria. Therefore, it is very likely that these laminations were
preserved due to the absence of bioturbating organism rather than the
presence of anoxia. The absence of anoxia proposed here is supported
by similar investigations in Nanpanjiang Basin of Chaina (Krull et al.,
2004), in Abadeh section of Iran (Heydari et al., 2003; Korte et al.,
2004b), in several Tethyan sections (Kozur, 2007), and in Central
Mountains of Oman (Twitchett et al., 2004). The rare organic-rich
deposits at the PTB interval (i.e., Thomas et al., 2004) could have
formed in marginal basins, similar to the anoxic water columns of the
Black Sea or the Cariaco basins that occur adjacent to the fully
oxygenated modern oceans.
4.3. A change in the style of carbonate production
The uppermost Permian, pre-Event Horizon limestone presented
in this study is similar to all other limestones of Cambrian to Recent,
suggesting that they were produced by a carbonate factory that was
typical of the Phanerozoic Earth (Scholle et al., 1983; Schlager, 2005).
This means that carbonates were produced by biochemical processes
in the photic zone (Scholle et al., 1983; Schlager, 2005).
Deposition of carbonate-poor strata of the Transition Zone interval
may indicate a halt in carbonate production after the mass mortality. The
absence of carbonate production was most likely because of a decrease
in oceanic pH as will be discussed below. Full carbonate production
began in the Early Triassic. But, the type of carbonate generated was
different from those produced prior to the Event Horizon. The lowermost Triassic limestone in the Shahreza section and elsewhere globally
(see Pruss et al., 2006 for review) contain abundant evidence of
microbial activity as well as the growth of carbonates on the seafloor by
abiotic processes, both characteristics of the Precambrian Earth
(Grotzinger and Kasting, 1993). The increase in microbial activity cannot
be attributed to the absence of metazoan decline (see Riding, 2006). In
addition, spontaneous carbonate precipitation such as the growth of
cement on the seafloor requires a very high carbonate supersaturation of
seawater (Morse and Mackenzie, 1990). Therefore, we suggest that the
observed change in carbonate production from a Phanerozoic-type prior
to the Event Horizon to a Precambrian-type after the Event Horizon and
during the Early Triassic was caused by an increase in carbonate
saturation of seawater (see below).
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4.4. A variation in the original carbonate mineralogy
The uppermost Permian, pre-Event Horizon limestone consists of
small crystals (Fig. 3A), does not show recrystallization (Fig. 3A), and
contains low Sr concentrations (Fig. 2D). Such features have been
attributed to a limestone whose original mineralogy was calcite, a
stable carbonate mineral which resists recrystallization (Sandberg,
1983; Wilkinson et al., 1985; Hardie, 1996). In contrast, the post-Event
Horizon and the lowermost Triassic limestone including the crystal
fans consists of large crystals, exhibits total obliteration of original
fabrics by recrystallization (Fig. 3B), and has high and fluctuating Sr
compositions (Fig. 2D). These characteristics are attributed to a
limestone whose original mineralogy was aragonite, an unstable
carbonate mineral that experiences extensive dissolution after
deposition (Sandberg, 1983; Wilkinson et al., 1985; Hardie, 1996).
Therefore, a change in carbonate mineralogy was associated with the
end-Permian mass extinction. Calcite was the primary carbonate
mineral in pre-Event Horizon limestone; whereas, aragonite dominated the mineralogy of carbonates deposited after the Event Horizon
and during the Lower Triassic.
Petrographic and trace element results from the Chaotian and the
Julfa sections are not available to evaluate whether the same changes
occurred in these two sections. Such a systematic investigation has not
been performed for other localities either. However, numerous
descriptions point out that Early Triassic cements were originally
aragonitic (Woods et al., 1999; Pruss et al., 2005; Baud et al., 2005,
2007; Lehrmann et al., 2007; Woods et al., 2007).
This investigation confirms the result of Railsback and Anderson
(1987). However, broad and low resolution observations of carbonate
mineralogy and fluid inclusion studies have led some to suggest that
the Mississippian to the Triassic interval belonged to an aragonite sea
(Sandberg, 1983; Wilkinson et al., 1985; Hardie, 1996; Lowenstein et
al., 2001). The change in the original mineralogy of carbonates
proposed in this study could have been short-term perturbation
associated with the end-Permian event.
Most importantly, aragonite versus calcite mineralogy of marine
carbonates is also an indicative of changes in seawater composition
whose causal mechanism has been debated (Sandberg, 1983; Wilkinson
et al., 1985; Railsback and Anderson, 1987; Hardie, 1996). Some consider
Mg/Ca ratio of seawater to play the dominant role (Sandberg, 1983;
Hardie, 1996; Dickson, 2002); others suggest that carbonate saturation
of seawater is the primary control on carbonate mineralogy (Heydari
and Moore, 1994; Locklair and Lerman, 2005). Therefore, our observed
changes in original mineralogy of precipitating carbonates indicate
a change in carbonate saturation of seawater accompanied the endPermian mass extinction.
4.5. A decrease in δ13C composition of seawater
Marine carbonates of the Shahreza and Julfa sections shows 4–5‰
PDB decline in δ13C composition during the transition from the
Permian to the Triassic. No δ13C data is currently available for the
Chaotian section. Similar or even larger values than presented here
have been observed globally (Baud et al., 1989; Musashi et al., 2001;
Dolenec et al., 2001, 2004; Krull et al., 2004; Riccardi et al., 2006; Haas
et al., 2007; Horacek et al., 2007a,b). Decreases have also been reported
in δ13C composition of marine organic carbon (Wang et al., 1994; Krull
et al., 2004; Riccardi et al., 2006). Such low δ13C compositions indicate
addition of carbon with low δ13C values to the latest Permian ocean.
Interpretations differ about how to achieve the low δ13C values at the
PTB (see Corsetti et al., 2005 for an excellent review).
Holser et al. (1989) suggested that δ13C decrease is related to
oxidation of organic matter during a sea-level fall. This interpretation
as well as others which advocate exchange between organic and
inorganic carbon reservoirs (i.e., Broecker and Peacock, 1999) is now
considered unlikely because of the slow nature of such events to
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account for the sharp decline in δ13C compositions observed at the
PTB (Erwin, 1993).
Hsu and McKenzie (1990) interpreted the declining δ13C compositions of the PTB as a signature of a respiring ocean, where bacterial
activity produced a vertical δ13C gradient that was the reverse of the
one produced by the biological pump in the modern ocean. This
interpretation has also been rejected because bacterial respiration
cannot be more than 100% efficient, which is necessary to cancel the
effect of productivity (Holser and Magaritz, 1992).
Gruszczyński et al. (1989, 1992) related the declining δ13C
compositions of the PTB section in West Spitsbergen to mixing of
water masses of a stratified ocean consisting of an anoxic bottom
water mass and an oxygenated surface water mass. However, the
isotopic compositions of West Spitsbergen section are now considered
diagenetic (Mii et al., 1997).
The concept of a stratified water column that was developed by
Gruszczyński et al. (1989, 1992) on data which were diagenetic in origin
has become the “model of choice” to explain not only the δ13C compositions but other characteristics of the PTB event including lithofacies,
cementation, and the mass extinction itself (see Kajiwara et al., 1994;
Knoll et al., 1996; Isozaki,1997; Woods et al., 1999; Heydari et al., 2000;
Kump et al., 2005; Riccardi et al., 2006). The problem is that the
existence of an anoxic water mass that extended to shallow waters is
not supported by sedimentological record as discussed in previous
paragraphs. Furthermore, the mixing of its water masses cannot
account for many features of the PTB including up to 8‰ PDB shift in
δ13C across this boundary (see Heydari and Hassanzadeh, 2003; Krull
et al., 2004).
Several hypotheses relate the decrease in δ13C to CH4. Krull and
Retallack (2000) and Krull et al. (2004) suggested that the decline in
δ13C is related to the release of CH4 primarily from high latitude polar
regions. Although polar region contributed to the decrease in δ13C but
it is unlikely that all of the decline could have been due to release
of CH4 from the polar region. Ryskin (2003) suggested that CH4
accumulated in the bottom waters of a stagnant ocean which was
erupted during the end-Permian time. No physical evidence of a
stagnant ocean exists, nor do we have any sedimentological record of
such a gigantic oceanic gas eruption. In addition, basic problems exist
with the amount of CH4 which actually can be accumulated in the
deep ocean (see Dickens, 2004). Heydari and Hassanzadeh (2003)
proposed that δ13C composition of the PTB strata is related to the
dissociation of deep-water marine gas hydrates and the release of CH4
to the ocean. Microbial oxidation of CH4 in the water column resulted
in the release of CO2 with low δ13C composition leading to low carbon
isotopic composition of marine organic and inorganic carbon.
Recently, Retallack et al. (2006) have proposed that coal metamorphism associated with intrusion of igneous dikes was responsible for
these low values.
A comprehensive modeling approach was attempted by Payne and
Kump (2007) to resolve the cause of the δ13C perturbations at the PTB
as well as those of the Early Triassic. At the end, these authors
concluded that a combination of volcanic-derived CO2 and organic
carbon would explain positive and negative fluctuations in δ13C
composition of marine carbonates of this time interval. Except for
the PTB perturbation, we do not have any evidence of volcanic activity
during the Lower Triassic, however. Furthermore, the terrestrial
organic carbon delivery to the ocean was very low during this time
interval (Berner, 2005). In addition, the input data of Payne and Kump
(2007) may be unrealistic. For example, in their single perturbation
approach, these authors used 3 × 1018 mol C of volcanic origin. This
value amounts to 36 × 1018 g carbon, which is more than three times
greater than value calculated by Kamo et al. (2003) for the Siberian
Traps volcanism (about 11 × 1018 g of carbon). In fact, volcanic-derived
carbon of Payne and Kump (2007) is more than all of short-term
exchange carbon of modern Earth (34 × 1018 g of carbon, see Dickens et
al., 1995). Yet, even this amount of volcanic-derived carbon produce
less than 1‰ PDB negative perturbation (Payne and Kump, 2007, their
Fig. 3). The best result of the modeling approach by Payne and Kump
(2007) which fits the observed δ13C variations at the PTB is a single
perturbation by gas-hydrated derived carbon, a case which these
authors eliminated as possibility.
4.6. Short comings of previous mass extinction models
No physical evidence (shocked quartz, spherules, microtektites,
tsunami deposits) have been found at or adjacent to the PTB to justify
a bolide impact (Heydari and Hassanzadeh, 2003). In addition, the
proposed geochemical indicators of extraterrestrial cause of the endPermian mass extinction have been adequately questioned (Farley and
Mukhopadhyay, 2001; Isozaki, 2001; Koeberl et al., 2004; Farly et al.,
2005). The recently proposed “Bedout impact crater” appears to be a
basement high (Müller et al., 2005).
As discussed in previous paragraphs, physical sedimentology and
paleontological evidence point to oxygenation of open marine environment during the latest Permian in China, Iran, Oman, and other sites in
the Tethyan region (Heydari et al., 2003; Krull et al., 2004; Korte et al.,
2004b; Twitchett et al., 2004; Chen et al., 2006a,b; Farabegoli et al., 2007;
Kozur, 2007). Anoxia could have exited in deep-water depressions as
well as coastal lagoons, while the open ocean remained oxygenated. This
is analogous to co-occurrence of the anoxic water column in the Black
Sea and the Cariaco Basin adjacent to the oxygenated modern ocean.
Therefore, varieties of anoxia-related mass extinction scenarios (Wignall
and Hallam, 1992; Knoll et al., 1996; Isozaki, 1997; Kump et al., 2005;
Riccardi et al., 2006) do not fit sedimentological record. Furthermore, the
recent suggestions that anoxia was extended all the way to the photic
zone during the end-Permian mass extinction (Grice et al., 2005; Kump
et al., 2005; Hays et al., 2007) and H2S was bubbling out of water does not
fit the physical sedimentological features of open marine strata in the
Chaotian, Julfa, and Shahreza sections. Furthermore, an ocean filled to
the rim with dissolved H2S is hard to justify in the presence of an
oxygenated atmosphere. Biomarkers used to suggest photic zone anoxia
may not be good oxygen indicators (Kirschvink, 2007). Lastly, numerical
model simulating appears to be of little help in resolving the ocean
circulation and water column oxygenation of the end-Permian ocean.
Excellent investigation by Winguth and Maier-Reimer (2005) demonstrated that various outcomes are possible depending on the underlying
assumptions.
Another commonly used end-Permian mass extinction model is
the eruption of the Siberian Traps volcanism only because of its
temporal co-existence of the two (Renne and Basu, 1991; Wignall,
2001; Kamo et al., 2003; Grard et al., 2005). The supporters of this
interpretation propose that the CO2 produced by volcanism triggered
a series of reactions that caused all of the events that characterize the
transition from the Permian to the Triassic (Renne and Basu, 1991;
Wignall, 2001; Kamo et al., 2003; Grard et al., 2005). The problem is
that continental flood basalts generate too small of a volume of CO2 to
cause any global change (Self et al., 2005). In fact, flood basalt CO2 is
not produced at once, rather it is generated by many eruptions each
100 s–1000 s of years apart (Self et al., 2005). Additionally, volcanic
rocks are a large sink for atmospheric CO2 through chemical weathering (Dessert et al., 2003; Self et al., 2005). Therefore, The CO2 that is
emitted from each flow is either consumed during the weathering of
volcanic rocks themselves, or re-equilibrated by the atmosphere (Self
et al., 2005). Furthermore, the δ13C composition of volcanic-derived
CO2 (−7‰ PDB) is not low enough to produce the observed declines in
δ13C values of marine and terrestrial system (Dickens et al., 1995) as
was shown convincingly by Payne and Kump (2007). The large
decrease in δ13C requires the involvement of CH4 (Berner, 2002).
The temporal correlation of the Siberian Traps volcanism and the
mass extinction does not imply a cause and effect relationship between
these two events as has been implied (Renne and Basu, 1991; Wignall,
2001; Kamo et al., 2003; Grard et al., 2005). Flood basalt eruption on
E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
land did not cause the mass extinction in the ocean. Rather, we suggest
that the Siberian Traps flood basalt and the end-Permian mass
extinction had a common origin: they were both the result of processes
associated with an active mantle plume (sea below).
A variation of volcanic eruption-related mass extinction has
recently been proposed. According to this model, the end-Permian
mass extinction was related to explosive gas-rich felsic volcanism, not
the continental flood basalt (Morgan et al., 2004; Isozaki et al., 2007b).
There may have been some explosive volcanism associated with the
PTB, but it is unclear how such a local, short duration event could
produce such long-lasting global change in sedimentological, mineralogical, and geochemical signals that occur at the PTB.
From the CH4-related mass extinction models (Krull and Retallack,
2000; Krull et al., 2004; Ryskin, 2003; Heydari and Hassanzadeh,
2003), the mechanism by Heydari and Hassanzadeh (2003) who
proposed that a big burst of gas-hydrated derived CH4 into the ocean
near the end of the Permian produced a major change in seawater
chemistry is the most likely. This model seems to simultaneously
explain all of the observed physical and geochemical observations at
the end-Permian mass extinction.
4.7. The proposed kill mechanism
It was demonstrated that the end-Permian biological devastation was
characterized by changes in the style of carbonate production, a variation
in original carbonate mineralogy, and decreases in δ13C composition of
marine carbonates. Such features should be explained in the context of a
model that also addresses other events which are known to symbolize
this interval of Earth's history, including the following:
(1) The mass killing in marine and terrestrial environments
(Retallack, 1995, 1999; Twitchett et al., 2001; Payne et al., 2004;
Erwin, 2006),
(2) About 4–8‰ PDB decline in δ13C values of organic and inorganic
marine carbon (Baud et al., 1989; Wang et al., 1994; Krull et al.,
2004; Riccardi et al., 2006),
155
(3) Up to 10‰ PDB decrease in δ13C composition of terrestrial carbon
(Thackeray et al., 1990; Morante 1996; Krull and Retallack, 2000;
De Wit et al., 2002),
(4) Precipitation of synsedimentary features which had not
occurred since the Precambrian time (Sano and Kakashima,
1997; Woods et al., 1999; Kershaw et al., 1999; Lehrmann,
1999; Lehrman et al., 2003; Pruss and Bottjer, 2004, 2005;
Baud et al., 2005; Pruss et al., 2005, 2006; Woods et al., 2007;
Baud et al., 2007),
(5) Severe global warming which resulted in the growth of forest in
high latitude regions (Taylor et al., 1992; Retallack, 1999),
(6) The eruption of the Siberian Traps flood basalt (Renne and Basu,
1991; Reichow et al., 2002; Kamo et al., 2003; Mundil et al.,
2004),
(7) Abnormal pollen grains (Foster and Afonin, 2005).
(9) Increased chemical weathering (Martin and Macdougall, 1995;
Sheldon, 2006), and
(10) Enhanced soil erosion (Stephenson et al., 2005).
4.7.1. Mass killing in the ocean
The key in finding a mechanism that can simultaneously explain all
of the changes that occurred at the PTB lies in determining the origin of
the large declines in δ13C compositions of organic and inorganic carbon
in marine and terrestrial environments (Baud et al., 1989; Thackeray
et al., 1990; Wang et al., 1994; Morante 1996; Krull and Retallack, 2000;
De Wit et al., 2002; Krull et al., 2004; Korte et al., 2004a,b; Riccardi
et al., 2006; Richoz, 2006; Farabegoli et al., 2007; Galfetti et al., 2007;
Haas et al., 2007; Horacek et al., 2007a,b). Various scenarios proposed
to explain the decline in δ13C compositions were discussed in previous
sections, but none are as capable as injection of gas hydrate-derived
CH4 into the ocean. The problem with this scenario has been as how to
destabilize gas hydrates (see Katz et al., 2001).
Dissociation of marine gas hydrates needs either an increase in
temperature or a decrease in pressure which can be accomplished by
several mechanisms such as a sea-level drop, global warming, and
submarine slumping (Nisbet, 1990; Kvenvolden, 1993; Dickens et al.,
Fig. 4. Injection of igneous dike swarms, possible high heat flow, and fracturing associated with an active mantle plume in the latest Permian caused massive release of CH4 primarily
from the dissociation of marine gas hydrates and secondarily from accelerated maturation of organic-rich sediments and leaking of petroleum reservoirs. Aerobic oxidation of CH4 in
the water column and the subsequent changes in seawater composition resulted in the marine mass extinction and the observed low δ13C compositions of marine carbon seen
globally. Gas-charged-oceans released large amounts of CH4 and CO2 into the atmosphere causing a hot climate which in turn caused the dissociation of polar gas hydrates releasing
additional CH4 into the atmosphere intensifying the global warming. Oxidation of CH4 in that atmosphere produced CO2 with δ13C values. Combined, these produced low δ13C values
seen in terrestrial carbon. Feeder dikes formed the Siberian Traps flood basalt. Cooling caused by volcanic-derived SO2 was either short lived, not yet recognized, or in-effective
against the global warming produced by CH4 and CO2. However, high global temperatures combined, acid rain, soil acidification, soil erosion and loss of habitat caused the terrestrial
mass extinction.
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1995; Dickens, 2000). Changes in the relative sea-level and atmospheric temperature are unlikely to trigger gas hydrate dissociation in
the deep ocean (Milkov and Sassen, 2003). Furthermore, there is no
evidence to substantiate a submarine slump origin of CH4 release at
this time.
We suggest that the dissociation of the end-Permian gas hydrates
was linked to the same process that caused the eruption of the
Siberian Traps volcanism: i.e., a mantle plume. In fact, the Siberian
Traps volcanism indicates that a mantle plume indeed existed at the
PTB, because flood basalts are a proxy for the occurrence of mantle
plumes (Coffin and Eldholm, 1994; Condie, 2004; Campbell, 2005).
Mantle plumes are columns of hot, solid material that originate in the
lower mantle, probably at the core–mantle boundary, and rise upward
forming a mushroom-type structure at the base of the lithosphere
causing many perturbations in the Earth's system (Coffin and Eldholm,
1994; Condie, 2004; Campbell, 2005; Lay et al., 2006).
Events attributed to mantle plumes such as injections of igneous dike
swarms, fracturing, and uplift (Coffin and Eldholm, 1994; Condie, 2004;
Campbell, 2005) caused the rapid release of large volumes CH4 from the
dissociation of marine gas hydrates (Figs. 4–5). This methane was
injected into the end-Permian ocean where it was microbially oxidized
in the oxygenated water, similar to what is taking place in the modern
ocean (Paull et al., 1995; Suess et al., 1999; Valentine et al., 2001):
CH4 þ 2O2 →CO2 þ 2H2 O
ð1Þ
CO2 produced from the reaction 1 had the potential to cause several
major perturbations in the oceanic–atmospheric system including:
(1) lowering the δ13C composition of the ocean, (2) changing the
carbonate saturation state of seawater, (3) increasing global temperature, and (4) causing the marine mass extinction.
4.7.1.1. Carbon isotope decrease. The δ13C composition of the modern
biogenic CH4 gas varies from −40 to −80‰ PDB (Borowski et al., 1996;
Sassen et al., 1999; Suess et al., 1999). A value of −60‰ PDB is commonly
accepted (Dickens et al., 1995). The low δ13C value of gas hydratederived methane was fully capable of producing the necessary shift to
low δ13C values of marine organic and inorganic carbon seen at the PTB.
We can use the mass balance equation of Dickens et al. (1995) to
calculate the amount of CH4 needed to produce the observed δ13C shift
at the PTB:
ðMTX þ MR Þ d13 CTX V ¼ ðMR Þ d13 CR þ ðMTX Þ d13 CTX
ð2Þ
Where, MTX and δ13CTX are mass and carbon isotope composition,
respectively, of total exchangeable carbon from exogenic reservoirs
(organic and inorganic); MR and δ13CR are mass and carbon isotope
values, respectively, of carbon transferred from an external reservoir
(gas hydrates). The δ13CTX' is the carbon isotope value of total carbon
reservoir after exchange has taken place (see Dickens et al., 1995).
In order to produce a −4‰ PDB decline in δ13C compositions of marine
carbonate at the PDB, about 1.2×1018 g or 1200 Gt (1 Gt=1015 g) of carbon
is needed. The global estimate on the total amount of carbon in gas
hydrates vary. Estimates range from 5000–10,000 Gt (see Kvenvolden,
1988; Dickens et al.,1995; Milkov et al., 2003). Assuming a value of 5000 Gt
Fig. 5. This chart summarizes simultaneous events that took place on land and in sea which led to biological devastations in marine and terrestrial environments and produced
sedimentological, mineralogical, and geochemical changes across the PTB.
E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
proposed by Buffett and Archer (2004), then, a −4‰ δ13C perturbation at
the PTB can be accomplished by the release of 25% of marine gas hydrate.
The 25% dissociation of marine gas hydrates is the maximum
needed. This is because of the availability of other sources of CH4 and
CO2 associated with the injection of igneous dike swarms. These
include CH4 and CO2 contributions from accelerated maturation of
organic-rich sediments similar to the Paleocene–Eocene boundary
(Svensen et al., 2004) and the release of these gases by fracturing of
petroleum reservoirs as in the modern Gulf of Mexico (Roberts and
Aharon, 1994; Sassen et al., 1999, 2004).
4.7.1.2. Carbonate saturation of the ocean. The CO2 that was produced
via the reaction 1 would have decrease pH of seawater and influenced
carbonate saturation of the end-Permian ocean. This is expressed
using CO2–carbonic acid system (Morse and MacKenzie, 1990):
CO2ðgÞ ↔CO2ðaqÞ
ð3Þ
CO2ðaqÞ þ H2 OðlÞ ↔H2 CO3ðaqÞ
ð4Þ
H2 CO3ðaqÞ ↔HCO−3 ðaqÞ− þ Hþ
ðaqÞ
ð5Þ
þ
HCO−3 ðaqÞ↔CO2−
3 ðaqÞ þ HðaqÞ
ð6Þ
Ac ¼
mHCO−3
þ
2mCO2−
3
ð7Þ
2−
Ca2þ
ðaqÞ þ CO3 ðaqÞ↔CaCO3ðsÞ
ð8Þ
Ω ¼ ðaCa2þ aCO2−
3 Þ=K
ð9Þ
Where a = activity, Ac = carbonate alkalinity, aq = aqueous, g = gas,
l = liquid, m = molality, K and Ω are the thermodynamic equilibrium
constant and the saturation state, respectively, of carbonate minerals
(aragonite or calcite). Even with no change in calcium content of
seawater, carbonate saturation will decrease because CO32− concentration declines with a lowering of the pH (Morse and MacKenzie, 1990).
Therefore, reactions 3–8 can be summarized as:
CaCO3ðsÞ þ CO2ðaqÞ þ H2 OðlÞ ↔2HCO−3 ðaqÞ þ Ca2þ
ðaqÞ
ð10Þ
As CO2 concentration increases the reaction proceeds to the right:
calcium carbonate will dissolve producing bicarbonate and calcium
ions. As CO2 deceases the reaction will proceed to the left: calcium
carbonate will precipitate.
A case example of the proposed scenario is the decrease in
carbonate saturation of the modern ocean due to the increase in
atmospheric CO2 (Caldeira and Wickett, 2003; Feely et al., 2004; Orr
et al., 2005). However, there are two differences between the modern
system and that of the end-Permian: (1) the source of CO2:
anthropogenic in the modern, gas hydrate at the end-Permian time,
and (2) point of origin of the CO2: from the atmosphere to the ocean in
the modern, from the ocean to the atmosphere during the endPermian. In fact, in one of the earliest attempts to explain the secular
variation in the original mineralogy of ooids during the Phanerozoic,
Heydari and Moore (1994) suggested that changes in marine
carbonate mineralogy was most likely caused by variation in
carbonate saturation of seawater as a function of atmospheric CO2.
Heydari and Moore (1994) demonstrated that seawater becomes
undersaturated with respect to aragonite with a seven fold increase in
partial pressure of atmospheric CO2.
In the Permian, as in the modern ocean, carbonates were produced
biochemically within the photic zone (see Shinn et al., 1989; Tucker
and Wright, 1990; Schlager, 2005). This carbonate production appears
to have been terminated at the Event Horizon and carbonate
157
deposition was replaced by siliciclastic shale in the Shahreza and
Julfa sections and by marl in the Chaotian section (Fig. 1). The
cessation of carbonate deposition at the Event Horizon is attributed to
an increase in CO2 concentration of the end-Permian ocean. Two cases
of end-Permian submarine dissolution have already been reported
(Heydari et al., 2003; Payne et al., 2007). Therefore, the end-Permian
CO2-saturated ocean became acidic: the “oceanic acid bath” of Heydari
and Hassanzadeh (2003).
4.7.1.3. A warm ocean. The end-Permian gas-charged ocean released
massive amount of CH4 and CO2 gases into the atmosphere producing
a very hot climate (the end-Permian Inferno) (Figs. 4–5). This hot
condition is supported by the migration of calcareous algae to the
Boreal region (Wignall et al., 1998), forested polar regions (Taylor et al.,
1992; Retallack, 1995), and the formation of paleosols at high latitudes
(Retallack, 1999; Krull and Retallack, 2000), and climate simulation
models (Kiel and Shields, 2005). This hot climate also produced a
warm ocean. In fact, δ18O composition of unaltered calcite of the
uppermost Permian and the lowermost Triassic shows one of the
highest temperatures of tropical settings (Veizer et al., 2000). Some
studies even suggest temperatures as high as 60 °C on the basis of very
low oxygen isotope compositions of cherts, and chert–phosphate pairs
(Karhu and Epstein, 1986; Railsback and Anderson, 1987).
4.7.1.4. Death in the ocean. The way organisms died is perhaps the least
understood aspect of the end-Permian mass extinction. Those who
advocate anoxia for this time interval also imply death due to the lack
of oxygen (Weidlich et al., 2003) or worse yet the abundance of H2S all
the way to the photic zone (Kump et al., 2005). The most direct
analysis of death mechanism was conducted by Knoll et al. (1996) who
concluded the end-Permian biological devastation was due to high
CO2 concentration (hypercapnia). However, even the elaborate study
of Knoll et al. (1996) seems to have been very simplistic. This is
because the main conclusion of Knoll et al. (1996) that organisms with
weak internal circulation and low metabolic rates were more sensitive
to hypercapnia is rejected based on detailed study of biological
behavior under adverse conditions by Pörtner et al. (2004). Furthermore, investigations by Pörtner et al. (2004) Pörtner and Langenbuch,
(2005) suggest that the issue of how organisms behave under stressful
condition is far more complex than generally realized.
Our investigation provides a very general environmental condition
which led the death in the ocean. To summarize, the dissociation of
marine gas hydrates and the release of CH4 into to the ocean resulted
in seawater that was CO2-charged, acidic ocean that had high
concentrations of Ca2+ and HCO3−, and low CO32− values (Figs. 4–5).
We suggest that it was these changes in seawater composition
combined with warm seawater temperature the caused the mass
mortality in the ocean. The acidic pH of the seawater made
calcification of many marine biota very difficult (Kleypas et al., 1999;
Langdon et al., 2000; Riebesell et al., 2000; Riebesell, 2004; Andersson
et al., 2005). The combined effects of high temperature and CO2
concentration caused major physiological crisis in organisms including reduced metabolic rates, high sensitivity to environmental stress,
and hampered growth and reproduction (Clarke, 1993; Pörtner et al.,
2004; Pörtner and Langenbuch, 2005). Therefore, we propose that
physiological limitations imposed on organisms by a rapid change in
the chemical composition of seawater such as increased concentration
of CO2, low pH, low CO32− values, high Ca2+ and HCO3− concentrations,
and high temperatures that caused the end-Permian mass extinction.
4.8. Devastation on land
The release of CO2 and CH4 (both greenhouse gases) from the endPermian ocean into the atmosphere had several devastating consequences. (1) It produced such a global warming that polar regions
became forested (Taylor et al., 1992; Retallack, 1999). As indicated, this
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E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
global warming also raised seawater temperature intensifying oceanic
mass extinction that was triggered due to high dissolved CO2 of
seawater. (2) The global warming in turn caused the dissociation of
polar gas hydrates releasing additional CH4 into the atmosphere
which further fueled the global warming (Figs. 4–5). Perhaps, it was
the addition of this CH4 to the system which resulted in lower than
normal δ13C signatures in polar regions discovered by Krull and
Retallack (2000; Krull et al., 2004). (3) CH4 oxidation in the
atmosphere produced CO2 with low δ13C values which resulted in
the observed δ13C decline of terrestrial carbon (Thackeray et al.,
1990; Morante, 1996; Krull and Retallack, 2000; De Wit et al., 2002).
(4) The global warming as well as high CO2 concentration of the
atmosphere intensified chemical weathering as shown by enhanced
soil erosion, high delivery of fine particles to the ocean, and
increases in 87Sr/86Sr isotope composition of seawater (Martin and
Macdougall, 1995; Stephenson et al., 2005; Sheldon, 2006).
Contemporaneously, feeder dikes associated with the mantle
plume formed the Siberian Traps (Figs. 4–5). Volcanism was not the
cause or the trigger for the PTB mass extinction. It was just another
event associated with the end-Permian mantle plume. The CO2
emitted from the trap was neither sufficient to cause a global warming
(see Self et al., 2005) nor its isotopic value low enough to produce the
δ13C decline in marine and terrestrial carbon (see Payne and Kump,
2007). However, as demonstrated from modern volcanism (Wignall,
2001; Self et al., 2005), SO2 and aerosol emitted during the eruption of
the Siberian Traps flood basalt could have produced a climatic cooling
at the PTB that was either short lived, not yet recognized, or
insufficient to overcome the global warming imposed by CO2 and
CH4 content of the atmosphere.
The end-Permian inferno produced environmental stress on land
as shown by the abundance of abnormal pollen grains (Foster and
Afonin, 2005). Plant mass extinction was slow and shows regional
patterns (Rees, 2002; Goman'kov, 2005). Plant diversity decreased in
low latitude but increased in northern high latitudes during the PTB
interval (Rees, 2002). Therefore, the mass extinction on land can be
attributed to following events: (1) global warming produced by
increased atmospheric CO2, (2) acid rain that was generated by SO2
derived from the Siberian trap eruptions, (3) soil acidification caused
by increased atmospheric CO2 and SO2 (Oh and Richter, 2004), and (4)
loss of habitat.
4.9. Recovery in the ocean
Cessation of the mantle plume activity during Early Triassic
stopped the production of CH4 in the ocean. Therefore, the supply of
CO2 with low δ13C values to the seawater was terminated, as shown by
increases in δ13C compositions of marine carbon during the earliest
Triassic (Baud et al., 1989; Wang et al., 1994; Morante, 1996; Krull et al.,
2004). This decrease in dissolved CO2 led to an instantaneous increase
in carbonate saturation of seawater (the Soda Bath Ocean of Heydari
and Hassanzadeh, 2003) which had a profound effect on deposition of
marine carbonates (Morse and MacKenzie, 1990). Studies of Early
Triassic strata indicate that the crystal fans were aragonitic in
composition (Woods et al., 1999; Pruss et al., 2005; Baud et al., 2005,
2007; Lehrmann et al., 2007; Woods et al., 2007), suggesting that
aragonite, rather than calcite, became the primary carbonate mineral
precipitate. Massive seafloor cementation occurred and microbial
community flourished, both resulting in a Precambrian-type carbonate
production observed in the lowermost Triassic strata worldwide (see
Pruss et al., 2006).
4.10. Recovery on land
The flow of CO2 from the ocean into the atmosphere declined
because of the low dissolved CO2 of the seawater. Mild temperatures
and low atmospheric CO2 decreased chemical weathering and soil
erosion. Combined, these events ended the harsh atmospheric
condition and allowed plants to begin their recovery as shown by
palynological record (Foster and Afonin, 2005).
5. The dilemma of the marine sulfur isotope variation at the PTB
It has been known that the δ34S composition of the Late Permian
seawater was about +12‰ CDT (Canyon Diablo Troilite) and linearly
increased to as high as +30‰ CDT in the Early Triassic (Holser and
Kaplan, 1966; Strauss, 1997; Kampschulte and Strauss, 2004). This longterm change can be attributed to the exchange between the reservoirs of
oxidized sulfur (sulfates, high δ34S) and reduced sulfur (pyrite, low δ34S)
(Holser and Kaplan,1966; Strauss, 1997; Kampschulte and Strauss, 2004).
Removal of reduced sulfur through pyrite burial will increase δ34S
composition of seawater. Addition of reduced sulfur (pyrite weathering)
will lower the δ34S of ocean (Holser and Kaplan, 1966; Strauss, 1997;
Kampschulte and Strauss, 2004). Of course a variety of factors including
global tectonics, climate, sea-level change, and weathering rates will
influence this exchange (see Paytan et al., 1998).
Kaiho et al. (2001) documented as much as 25‰ CDT decline in
δ34S composition of carbonate associated sulfate (CAS) at the Event
Horizon of the Meishan section of China which they interpreted as a
sign of sulfur release from the mantle as a result of an asteroid impact.
Although this interpretation of δ34S has been questioned (see Koeberl
et al., 2002), but the investigation of the sulfur isotope variations as a
mean of resolving the end-Permian crisis has continued.
The result of Kaiho et al. (2001) and subsequent analyses raise
doubts about δ34S data collected. Firstly, δ34S values reported by Kaiho
et al. (2001) begin their decline from a value of +30‰ CDT (an unusual
value for the Late Permian which had the lowest δ34S of the
Phanerozoic, +12‰ CDT) and decline to approximately +5‰ CDT at
the Event Horizon. However, an exactly opposite trend was reported
by Newton et al. (2004) from the Siusi section of Italy where δ34S
composition of CAS begins an increase from a value of +10‰ CDT to a
value of about + 20‰ CDT at the Event Horizon. Even wilder
fluctuations were reported by Riccardi et al. (2006) from the Shangsi
section of China where CAS δ34S values begin their increase from a
value of −35‰ CDT to a value of +25‰ CDT at the Event Horizon, an
increase of about 60‰ CDT. To make the matter worse, Riccardi et al.
(2006) did not reproduce the results from the Shangsi section in
the Meishan section of china. Contradictory data have also been
demonstrated for δ34S of pyrite adjacent to PTB by Strauss (1997).
Therefore, until consistent and reliable δ34S compositions are
available, no reasonable interpretations can be made from sulfur
isotope data associated with the PTB interval.
6. Duration of the apocalypse
We consider the end-Permian mass extinction as a process
consisting of three distinct events: (1) the trigger, (2) the threshold,
and (3) the recovery. The trigger was the event that initiated the PTB
disturbance. The threshold was the point when the condition became
intolerable to organisms resulting in rapid mass mortality. The
recovery began when harsh environmental conditions ended.
We suggest that δ13C compositions of the Shahreza sections can be
used as a proxy to identify the trigger, the threshold, and the recovery
events (Fig. 6). The δ13C trend displays five major inflection points:
“A”, “B”, “C”, “D”, and “E” (Fig. 6). Point “A” occurs at about 9 m below
the Event Horizon and marks the point at which the decline in δ13C
values began (Fig. 6). It represents the start of the process which
eventually led to the end-Permian mass extinction: the trigger point
(Fig. 6). According to the kill mechanism of this study, the trigger
represent the time of gas hydrate dissociation and the injection of CH4
into the ocean.
No geochronological data are available from the Shahreza sections
to determine the time of the trigger. However, the red, nodular
E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
Fig. 6. Graph shows δ13C compositions of the bulk limestone at the Shahreza section of
Iran. Five inflections points are recognized. Point “A” (the trigger point) marks the start
of gas hydrate dissociation which led to the end-Permian mass extinction. The rate of
environmental deterioration as approximated by the rate of change in δ13C compositions was slow from “A” to “B”, but intensified from point “B” to point “C”. Point “C” (the
Event Horizon) marks the threshold when environmental deterioration reached
intolerable conditions and the mass mortality occurred. Point “D” is the paleontologically defined PTB and a failed attempt at the recovery. Point “E” marks the start of the
sustained recovery. Using a crude estimate form sedimentation rates, the uppermost
Permian to the lowermost Triassic process (from the trigger to recovery) could have
lasted 1.5 to 2 Myr (see text for details).
lithofacies in which point “A” occurs is 17–19 m thick and was
deposited during Dorashamian Stage (equivalent to Changhsing
Stage), believed to have lasted about 3 Myr (Bowring et al., 1998;
Heydari et al., 2003). This indicates a sediment accumulation rate of
about 6 m/Myr, or 167 Kyr per deposition of 1 m of sediment.
Therefore, end-Permian mass extinction was triggered at about
1.5 Myr before the Event Horizon or point “C” (Fig. 6).
The rate of environmental deterioration, approximated by the rate
of change in δ13C values, was very slow at first: about 0.5‰ PDB
change during the first 5 m of δ13C decline (from point “A” to “B”) or
during the initial 835 Kyr (Fig. 6). The rate of change in δ13C values,
and therefore the rate of environmental deterioration, increased
significantly from point “B” to point “C”: 1.5‰ PDB in 4 m of sediment
deposition or during the 668 Kyr before the Event Horizon or the point
“C” (Fig. 6). Our data indicate that the decline in δ13C value began and
ended entirely during deposition of a red, nodular wackestone
(lithofacies 1 in Fig. 1D) in deep, oxygenated waters. Therefore, the
killing process occurred entirely within an oxygenated environment.
Point “C” marks the Event Horizon or the time at which the
threshold was reached and organisms died rapidly (Fig. 6). According
to the kill mechanism of this study, this is the time in which the ocean
became CO2-charged and acidic disrupting physiological process of
marine organisms leading to rapid mass extinction. Therefore, it took
1.5 Myr (835 Kyr of slow change and 670 Kyr of rapid change) of
deterioration to reach the Event Horizon (Fig. 6). Unfortunately, the
duration of the threshold hold cannot be estimated from our data.
Although varies erratically, δ13C compositions continued to decline
from point “C” (the threshold) to point “D” (the paleontologically defined
159
PTB) indicating that the process that resulted in the end-Permian mass
extinction had remained for another 2 m of sedimentation (Fig. 6). The
δ13C value increased by about 1‰ PDB at point “D” which is considered to
be the first attempt at the recovery that failed. This failed recovery could
have been associated with a halt in the injection of CH4 into the ocean.
Sedimentation rate after the event horizon cannot be estimated because
of changes in the sediment type (from carbonate to shale), carbonate
mineralogy (from calcite to aragonite) occurred after the Event Horizon,
and submarine dissolution.
The δ13C composition began to decline after point “D” reaching its
lowest value at point “E” suggesting that the deterioration of the
environment continued for another 4 m after the paleontologically
defined PTB or point “D” (Fig. 6). Again, it is difficult to estimate the
rate of sedimentation from points “D” to “E” because there was a major
change in the nature of carbonate factory as indicated in previous
sections (from biochemical to microbial-abiotic).
The δ13C began a steady increase after point “E”: the point of
sustained recovery. Therefore, point “E” corresponds to the time at which
the dissociation of gas hydrates had ended and the injection of CH4 to the
ocean was terminated. In our analysis, the time interval between the
trigger (point “A”) and recovery (point “E”) is considered to be the
duration of the PTB process which could have lasted for at least 2 Myr.
This duration of the processes envisioned for the marine mass extinction
matches those considered for terrestrial mass mortality (see Goman'kov,
2005).
The three phases of the end-Permian mass extinction proposed by
Heydari and Hassanzadeh (2003) are shown on the δ13C compositional trend of the Shahreza section (Fig. 6). The Pardeess (paradise)
Phase signifying the equable environmental conditions that existed
prior to the point of trigger or point “A”. The Doozakh (inferno) Phase
indicating very hot and intolerable conditions occurred between
points “A” and “E”. The Rastaakheez (resurrection) Phase marking the
point rebuilding the environment began after point “E” (Fig. 6).
7. Conclusions
This study documents that the uppermost Permian limestones
were produced by a carbonate factory that was typical of the
Phanerozoic Earth. These limestones consisted of originally calcite
mineralogy with low and uniform Sr concentrations and display high
δ13C values. In contrast, the lowermost Triassic limestone was made
up of originally aragonite mineralogy, had high Sr values and low δ13C
signatures, and was produced by a Precambrian-type carbonate
production system via microbial and abiotic processes.
It is proposed that a mantle plume initiated a series of processes
which led the mass mortality on land and in sea and produced
sedimentological, mineralogical, and geochemical changes of the PTB
interval. Physiological limitations imposed on organisms by a rapid
change in the chemical composition of seawater such as increased
concentration of CO2, low pH, low CO32− value, high Ca2+ and HCO3−
concentrations, and high temperatures caused the end-Permian mass
extinction. Devastation on land can be attributed to global warming
produced by increased atmospheric CO2, acid rain that was generated
by SO2 derived from the Siberian trap eruption, soil acidification
caused by increased atmospheric CO2, and loss of habitat.
The end-Permian mass extinction was a process consisting of three
events: the trigger, the threshold, and the recovery. According to our
analysis, the PTB process (from the trigger to the recovery) could have
lasted for at least 2 Myr.
Several major mass extinctions of the Phanerozoic are temporally
accompanied by flood basalt eruption (Wignall, 2001). So far, this cooccurrence has been interpreted in a cause-and-effect relation: flood
basalts caused mass extinctions (i.e., Wignall, 2001). This study
proposes that these mass extinctions and their time correlative flood
basalt are parts of a much larger perturbation caused by mantle plume
activities. Without the presence of an active mantle plume neither
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E. Heydari et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 264 (2008) 147–162
would take place. In other words, changes in the near surface of Earth
are ultimately controlled by changes in the mantle. The results of this
study can potentially explain other association of continental flood
basalts and mass extinctions of the Phanerozoic. However, it is
particularly applicable to the one at the end-Guadalupian (Zhou et al.,
2002; Isozaki et al., 2007a).
Acknowledgements
This study was supported by Jackson State University, the
University of Tehran, Iran, and the University of Payame-Nour, Iran.
We are grateful to highly constructive comments from two anonymous reviewers.
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