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Australia: The Land Where Time Began |
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The volcanics
The Siberian Trap province has become a part of geological folklore
because of their immense size and they are often depicted to cover vast
areas (e.g. Officer & Page, 1996, p. 169). Flood basalts cover, in fact,
only 3.4 x 105 km2 of northwest Siberia, though
pyroclastics and intrusives, particularly sills, increase greatly the
local area of the total province to 1.5 x 106 km2
(Zolotukhin & Al’Mukhamedov, 1988). There is no doubt the original area
of the flood basalts was greater, though proposed figures of 5 x106
km2 have been proposed (Czamanske et
al., 1998; Kozur, 1998), are
probably exaggerated. According to Wignall estimates of their original
volume are even more speculative. The volume of the Siberian Traps that
have been assumed by many authors is 4 x 106 km3,
though only 4 x 105 km3 remains at present. A
thick pile of basalts from the Early Triassic was discovered in a
borehole from the western Siberian depression which suggests the
province may extend to the west beneath the sediments from the
Jurassic-Cretaceous of this major basin (Westphal et
al., 1998). Within the
province itself, however, it is suggested by the distribution of lavas
that they do not constitute a single continuous province, rather, the
amalgamation of several “subprovinces” (Mitchell et
al., 1994).
In western Siberia volcanism began with the formation of the Tuffaceous
Series (Sadovnikov & Orlova, (1993, 1998). These are predominantly
basaltic tuffs that underlie the flood basalts in most of the province
and they dominate the entire succession in the south of the region. It
has proved to be difficult to date the tuffs, but the discovery recently
of conchostracan faunas within interbedded strata, as well as their
comparison with similar faunas in Chinese sections that are well dated,
indicates that they span the middle part of the Late Permian (Kozur,
1998). Over wide areas flood basalt flows overlie the tuffs over wide
areas and it is suggested by recent radiometric dating efforts that the
onset of eruptions was coincident with the Permian-Triassic (P-Tr)
boundary. U-Pb dates were obtained from the ash band immediately below
the P-Tr boundary in South China (Claoué-Long et
al., 1991) of 251.2 ± 3.4 Ma.
Sanidine from the same ash band was analysed (Renne et
al., 1995) which obtained an
40Ar-39Ar age of 249.91 ± 1.52 Ma, which included
external error. When the same ash band was redated (Bowring et
al., 1998) they obtained a U-Pb
age of 241.4 ± 0.3 Ma. It is suggested by these dates that the P-Tr
boundary occurred between 250 and 251 Ma, with a greater likelihood for
the older end of this age range. Contrasting with this, initial 40Ar-39Ar
obtained from the Siberian Traps were substantially younger (Baksi &
Farrar, 1991), though the date (Renne & Basu’s, 1991) of 248.4 ± 0.3 Ma
(internal error only) 40Ar-39Ar from a basalt flow
near the base of the lava pile at Noril’sk was the first to indicate
that the onset of the eruption was close in age to the P-Tr boundary.
Redating of the standard (Fish Canyon sanidine) that was used to
calculate this age, produced an age of 250.0 ± 1.6 Ma and therefore an
even closer correspondence (Renne et
al., 1995), including all
uncertainties this date becomes 250.0 ± 2.3 Ma (Renne et
al., 1998). Within error this
is comparable to a U-Pb age of 251.2 ± 0.3 Ma that was obtained (Kamo et
al., 1996) from an intrusion
in the lower third of the Noril’sk lava pile. Samples were analysed (Venkatesan
et al., 1997) from throughout
the Noril’sk succession and obtained ages of 247.1 ± 1.9 Ma for the base
and 247.6 ± 2.5 Ma for the top of the lava pile which indicates a rapid
rate of eruption, possibly 1 Myr or less. A different standard was used
(Venkatesan et al., 1997) to
calculate their 40Ar-39Ar ages, which therefore
cannot be compared with the date of Renne et
al. though it is clear for
the Noril’sk succession, at least, the eruptions began around the P-Tr
boundary, persisting possibly for 1 million years or less. This
conclusion is, essentially, in accord with the biostratigraphic dating
that is available from the region, which suggests that the onset of
fissure eruptions occurred in the later part of the Dorashamian Stage,
which is the last stage of the Permian, also known as the Changxingian
Stage, slightly below the P-Tr boundary as defined by palynological data
(Sadovnikov & Orlova, 1993, 1998). A similar age assignment (Kozur,
1998) is suggested by conchostracan evidence. This age that is slightly
older is below the resolution that is available from radiometric dating.
Though the timing of eruptions in the Noril’sk region is constrained
reasonably well, it is important to remember that this area constitutes
only 7% of the total volume of the province (Venkatesan et
al., 1997), and so may not be
representative of the entire province. Up to 3 km of lavas infill a
pre-existing graben from the Late Permian in the Maimecha-Kotui area in
the NE of the province. The Maimecha-Kotui volcanics consists of a suite
of alkali-ultra basics that are chemically distinct, which appear to be
older than the Noril’sk succession, unlike the remainder of the flows in
the province, which are tholeiitic basalts. Thus an 40Ar-39Ag
plateau age was obtained (Basu et
al., 1995) of 253.0 ± 2.6 Ma from an olivine nephelinite in the
region. The same Fish Canyon sanidine standard was used to calibrate
this date, and the recalibrated age (Renne & Basu, 1991) was used to
recalculate the ages they reported (Renne & Basu, 1991). The age of
onset of Noril’sk fissure eruptions and Maimecha-Kotui eruptions can be
compared; they indicate that eruptions in the latter region occurred
0.7-5.0 Myr earlier at the 95% confidence level (Venkatesan et
al., 1997). It is clear that
much radiometric dating is required from elsewhere in the province.
Noril’sk provides much of the knowledge of the nature of the Siberian
Traps. In this area individual flows are small by the standards of
CFBPs, rarely exceeding a few 10s of metres in thickness and an extent
of a few 10s of km (Sharma, 1997). The eruption history therefore
consisted of many small flows of small volume, typically a flow of
10,000 m3 every 10,000 years. An unusually high proportion
(10%) of basaltic pyroclastics is contained in Noril’sk lava pile. In
total, there are approximately 30 beds, which range from a few 10s of
centimetres to more than 100 m thick (Venkatesan et
al., 1997). An important
component of the Maimecha-Kotui volcanics as well as in the succession
to the south of Putorana is tuffs. Their intrinsic cratonic setting is
also noteworthy. With the exception of the small Columbia River
Province, all other provinces occur at locations that are marked by
subsequent continental rifting, or in the case of the case of the
Emeishan and Panjal Volcanics, former sites. This relationship is
ascribed by many authors to a plume-related mechanism of continental
breakup, as noted above (Courtillot et
al., 1999). Some of the
predictions of the plume model, however, are not present in the Siberian
traps, notably the absence of pre-eruption doming (Czamanske et
al., 1998). It is suggested
by plume modelling of the size required to generate the Siberian traps
that up to 4 km of uplift should mark the arrival of the plume at the
base of the lithosphere (Farnetani and Richards, 1994). In western
Siberia, and in both the Noril’sk and Maimecha-Kotui areas, a
pre-existing graben topography in infilled by lava, which indicates that
rifting, and not uplift, predated the eruptions (Zorn & Vladimirov,
1989; Czamanske et al.,
1998). Nonetheless, the preferred option in most studies of the Siberian
Traps remains the plume model (Renne & Basu, 191; Veevers et
al., 1994; Conaghan et
al., 1994; Coffin & Eldholm,
1994; Sharma, 1997).
Extinction mechanisms
The significance of the Siberian Traps, whatever their origin, is well
established due to the central role played by them in the majority of
current models for the end-Permian mass extinction event. Attention has
been focused by the higher proportion of tuffs (relative to many other
CFBPs) in the Siberian Traps on the likely cooling effects of volcanic
dust and sulphate aerosols (Campbell et
al., 1992; Conaghan et
al., 1994; Renne et
al., 1992; Kamo et
al., 1996; Kozur, 1998). It
was therefore proposed (Campbell et
al., 1992) that cooling of
sufficient to cause an intense glaciation of the duration of the
Siberian Traps eruption – a period of 600 ka in their estimation. This,
in turn, is suggested to have been the cause of the major eustatic sea
level fall at the end of the Permian that has been widely reported (e.g.
Holser & Magaritz, 1987). It was suggested (Erwin, 1993), alternatively,
that broad uplift that centred on western Siberia, immediately before
the onset of volcanicity in the region, may have been recorded by the
fall. As noted above, however, this alternative is made untenable by the
lack of evidence for uplift in the region (Kamo et
al., 1996; Czamanske et
al., 1998). There is an equal
lack of evidence for end Permian glaciation, though it was suggested
(Campbell et al., 1992) that
it may have been too brief to have left physical evidence. Reassessment
of the latest Permian and earliest Triassic changes in sea level that
were recently carried out suggests that the absence of such evidence may
not be so puzzling as rapid sea level rise, not fall, is seen in most
sections at this time (Hallam & Wignall, 1999). Hitherto, much of the
evidence for a fall of sea level in the latest Permian has been based on
the absence of diagnostic biostratigraphic markers from the latest
Permian, though it was revealed by sequence stratigraphic analysis that
the latest Permian-earliest Triassic interval was marked by a phase of
rapid coastal onlap (Haq et al.,
1987; Wignall & Hallam, 1993). It has been speculated (Hallam, 1999)
that a major pulse of intracratonic volcanism may be reflected by the
sea level rise. The same effect has been proposed for some changes in
sea level that occurred in the Cretaceous (see below), though this
proposition has been made difficult to judge
for the P-Tr sea level changes by the absence of any pre-Jurassic
oceanic crust.
If glaciation can be discounted as a factor in the mass extinction event
in the end-Permian, the same cannot be said for global cooling. The
mechanism was favoured as the main cause of the mass extinction (in
particular Kozur, 1998). The principal evidence he based this on was
from the changing species of conodont thst were involved in distribution
during the crisis. In many sections of the Tethyan P-Tr boundary, the
appearance of
Clarkina carinata, a
species that has been interpreted to have a preference for cool water,
marks the extinction interval (Kozur, 1998). The temperature preferences
of
Clarkina species are,
however, by no means clear and it noteworthy that other species of
Clarkina, such as the
C. subcarinata which is
closely related, have been inferred to prefer warm water (Kozur, 1998).
Wignall suggests it needs to be considered that an alternative
possibility, that
C. carinata was a deep
benthic form that expanded its range during the deepening that was
caused by the sea level rise of the end-Permian. Contrasting with the
conodont evidence, contemporaneous changes in the terrestrial flora
suggested that forms that were cold adapted faired particularly badly
during the end-Permian crisis. In the high southern palaeolatitudes the
eradication of the
Glossopteris flora is a
particularly noteworthy feature of the event (Retallack, 1995).
Extinctions among the Boreal marine invertebrates in Spitzbergen were
also at least as severe as those from the lower latitude sections of
Tethys (Wignall et al.,
1998).
Acid rain, the other climatic consequence of SO2 eruptions,
has also been postulated as a cause of one of the more intriguing
aspects of the extinction interval: the proliferation of fungi. It was
revealed by palynological samples from the P-Tr sections around the
world an increase in fungal spores in the latest Permian times, in
particular in low palaeolatitudes where assemblages are dominated by
them (Eshet, 1992). This has been attributed to this degradation, by
volcanogenic acid rain, of floral ecosystems resulting in proliferation
of fungi on the ample decaying vegetation. An alternative origin was
proposed (Hallam & Wignall, 1997) linked to the extinction of many
insect orders at the end of the Permian (Labandeira & Sepkoski, 1993).
At present insects destroy a large proportion of terrestrial flora, as
they probably did in the Permian, but fungi may respond to the increased
availability of decaying vegetation in the absence of insects. According
to Wignall a factor that requires consideration (Wignall et
al., 1996; Kozur, 1998), is
the possibility that the fungal spores event records the proliferation
of marine rather than terrestrial fungi, and this weakens further the
link between the fungal spike and acid rain. The most obvious climatic signal that is associated with the mass extinction event of the end-Permian is a major phase of global warming. The evidence includes calcareous algae migrating to the Boreal latitudes (Wignall et al., 1998), the preferential loss of high latitude floras noted above (Retallack, 1995), and palaeosols developing that are typical of latitudes of 50o or lower at palaeolatitudes as high as 80o (Retallack, 1999; Retallack & Krull, 1999). Equatorial temperatures are suggested by oxygen isotope data to possibly have risen by as much as 6oC at the P-Tr boundary, though such inferences need to assume the fluctuations of salinity can be disregarded (Holser et al., 1989). Strontium isotope ratios in marine carbonates and phosphates also indicate an increase of CO2 around this time. After the low point at the end of the Middle Permian, the ratios of 37Ar/36Ar began to rise at an increasingly rapid rate, reaching a high point at the end to the Early Triassic (Martin & McDougal, 1995). Increased continental weathering in an atmosphere that was increasingly rich in CO2 may be reflected by this trend (Martin & Macdougall, 1995) and/or enhanced leaching of the volcanogenic acid rain (Conaghan et al., 1994). Alternatively, the enhanced continental erosion during major sea level fall (Holser & Magaritz, 1987), does not account for the lack of correlation between the eustatic sea level curve and the isotopic trends (Hallam & Wignall, 1997)
as reflected by the Sr isotope curve.
Wignall suggests that global warming may also have been the cause of the
rapid development of marine anoxia as seen in many shelf sections from
the latest Permian (Wignall & Twitchett, 1996). In the deep water
pelagic chert sections from the accreted terranes of Japan at the end of
the Permian this “superanoxic” event first developed (Isozaki, 1997),
though it is the expansion of oxygen-poor conditions into shallow waters
in the latest Permian which coincides with the marine mass extinction
event (Wignall & Hallam, 1992, 1996; Wignall et
al., 1995). The decline of
the equator-to-pole temperature gradient and the consequent increase in
oceanic circulation, as well as the lower solubility of oxygen in warmer
waters, are the 2 effects of warming that may have been responsible for
marine anoxia (Wignall & Twitchett, 1996). Though the ultimate cause of
the elevated atmospheric concentrations of CO2 required to
generate the warming is not known. The preserved source of many is
volcanic CO2 that was derived from the Siberian Traps (e.g.
Campbell et al., 1992;
Conaghan et al., 1994;
Veevers et al., 1994; Wignall
& Twitchett, 1996), though contributions from gas hydrates (Erwin, 1993;
Morante, 1996), and oxidation of coals from Gondwana (Faure et
al., 1995), may also have
been important.
Carbon isotopic trends
According to Wignall in marine as well as terrestrial sections during
the P-Tr interval dramatic δ13C fluctuations may hold a clue
to the cause of changes of atmospheric CO2. Values of δ13C
gradually decreased in marine carbonates through the Dorashamian Stage
prior to a rapid negative excursion, of 4-5‰ magnitude, began just below
the P-Tr boundary (Holser et al.,
1989, 1991; Xu & Yan, 1993; Morante et
al., 1994; Morante, 1996). A
brief influx of isotopically light carbon is indicated by this negative
spike into the ocean-atmosphere system around the time of the
end-Permian mass extinction event. A potential source is volcanogenic CO2,
though its δ13C value of -5‰ (McLean, 1995) is probably not
sufficiently light to achieve the swing that is observed. This becomes
readily appreciated when the amount of C in surficial reservoirs, which
is estimated to be around 5 x 109 g C (Berner, 1999), is
compared with the volumes of CO2 that has been estimated
during the formation of CFBPs. It has been calculated, based on
measurements from Hawaiian eruptions (McCartney et
al., 1990), that 5 x 1012
g of C is emitted by 1 km3 of basalt. Therefore, if it is
assumed that a high value of 2 x 106 km3 of basalt
was present originally in the Siberian Traps, and then 1 x 1019
g of CO2 C may have been released during the eruption. This
is enough to cause only roughly 20% of the isotopic swing that is
observed in the latest Permian. It is clear that there are many
assumptions implicit in this reasoning, not least of which is the -5%
isotopic composition of volcanic CO2. The debate on the
ultimate origin of CFBPs, in this context, is apposite to this
assumption. It was suggested (Anderson, 1994, 1999) that rather than
being sourced from deep mantle plumes, CFBPs may result from shallow
mantle enriched in recycled lithosphere melting. If this layer, that is
termed the perisphere, is enriched in isotopically light, subducted
organic C then the δ13C of the volcanogenic CO2
that was emitted from the source may be considerably lighter than the
-5‰ value. The isotope excursion that occurred at the end-Permian could
have been caused by the eruption being about 1 x 1019 g C
with a δ13C value of -20‰. Before these speculations are
accepted it would be salient, however, to recall that much, if not most
of the Siberian traps province was erupted in the early Triassic
after the δ13C
excursion. At the latest Permian the sudden influx of light C into the
atmosphere and oceans has been attributed to a near-total collapse of
primary productivity (Holser & Margaritz, 1992; Margaritz et
al, 1992; Wang et
al., 1994). Once again,
however, consideration of the masses involved highlights that are not
likely to cause more than a small negative shift. Therefore, the modern
biomass, which amounts to 8.3 x 1017g C, and which has an
isotopic composition of -20‰, if it was added to the 5 x 1019
g C that is present in the inorganic C component of the ocean-atmosphere
system, would clearly have little impact on values of δ13C.
It has been appreciated in recent years that one of the major
repositories of C is isotopically light C which occurs in the form of
methane hydrates (clathrates) that are buried at shallow depths beneath
cold and/deep seas (dickens et al.,
1997). Isotopically, hydrates are very light (δ13C = -65‰),
and there may be as much as 1 x 1019g of C present in the
hydrates beneath modern seas. The release of only 10% of this material
into the atmosphere as methane would be enough to cause the δ13C
shift that is observed across the P-Tr boundary (Erwin, 1993; Browning
et al., 1998). The magnitude
of the end-Permian δ13C excursion clearly renders its origin
problematic. The eruption of vast volumes of CO2 and the
shutdown of productivity may have occurred at this time, though these
events can account for only a fraction of the C isotope changes that are
observed. The warming effect of the release of CO2 may,
however, have triggered the disassociation of huge amounts of methane
hydrates, thereby producing the δ13C excursion, exacerbating
the warming trend. There is substantial geological and palaeontological
evidence for conditions that were globally warm in the Early Triassic,
as noted above. There have been proposals for similar “runaway
greenhouse” scenarios for the Cenomanian-Turonian interval of the
Cretaceous as well as in the Palaeocene (see below) with flood basalt
volcanism implicated in both cases.
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Author: M.H.Monroe Email: admin@austhrutime.com Sources & Further reading |