The Geochemistry of Mass Extinction
LR Kump, The Pennsylvania State University, University Park, PA, USA ã 2014 Elsevier Ltd. All rights reserved.
6.12.1 Introduction 6.12.2 Isotopic Records of the Major Mass Extinctions 220.127.116.11 Carbon Isotopic Record 18.104.22.168 Oxygen Isotope Record 22.214.171.124 Sulfur Isotopic Record 126.96.36.199 Strontium Isotope Record 6.12.3 Interpreting the Geochemical Records of Mass Extinction 188.8.131.52 Late Ordovician 184.108.40.206 Late Devonian 220.127.116.11 End-Permian 18.104.22.168 End-Triassic 22.214.171.124 Cretaceous–Paleogene 6.12.4 Summary with Extensions Acknowledgments References
Glossary Biologic pump The transfer of carbon and nutrients from the surface ocean to the deep by their passive or active uptake during phytoplankton growth and settling of organic matter and decomposition in the deep ocean. Calcifier An organism that builds its skeleton out of calcium carbonate. Chemocline A vertically oriented zone in a body of water where the chemistry of the water changes markedly, often coincides with a thermocline or density transition (pycnocline). Eutrophication The buildup of nutrients in a body of water, generally accompanied by an intensification of biologic productivity but sometimes with a decline in biodiversity.
The course of biologic evolution is inextricably linked to that of the environment through an intricate network of feedbacks that span all scales of space and time. Disruptions to the environment have biologic consequences, and vice versa. Fossils provide the prima facie evidence for biotic disruptions: catastrophic losses of global biodiversity at various times in the Phanerozoic. However, the forensic evidence for the causes and environmental consequences of these mass extinctions resides primarily in the geochemical and isotopic composition of sedimentary rocks deposited during the extinction intervals. Thus, advancement in the understanding of mass extinctions requires detailed knowledge obtained from both paleontological and geochemical records. This article reviews the state of knowledge concerning the geochemistry of the ‘big five’ extinctions of the Phanerozoic
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Euxinic A water body is said to be ‘euxinic’ if aqueous hydrogen sulfide is detectable. Negative isotope excursion A shift toward lower values of the standardized ratio of the isotopes from a relatively stable preceding baseline, generally followed by a return to that baseline. Positive isotope excursion A shift toward higher values of the standardized ratio of the isotopes from a relatively stable preceding baseline, generally followed by a return to that baseline. Stomatal density A measure of the areal abundance of stomata (pores) on a leaf surface, shown to be sensitive to the atmospheric carbon dioxide partial pressure.
(e.g., Sepkoski, 1993): the Late Ordovician (Hirnantian, 440 Ma), the Late Devonian (an extended or multiple event with its apex at the Frasnian–Famennian (F–F) boundary, 367 Ma), the end-Permian (end-P, 251 Ma), the end-Triassic (end-T, 200 Ma), and the Cretaceous–Paleogene (K–P, 65 Ma). The focus on the big five is a matter of convenience, as there is a continuum in extinction rates from ‘background’ to ‘mass extinction’ (Alroy, 2008). And although much of the literature on extinctions centers on the causes and extents of biodiversity loss, in recent years, paleontologists have begun to focus on recoveries (many papers in Alroy, 2008; Erwin, 2001; Hart, 1996; Kirchner and Weil, 2000). To the extent that the duration of the recovery interval may reflect a slow relaxation of the environment from perturbation, analysis of the geochemical record of recovery is an integral part of this effort. In interpreting the geochemical and biologic records of recovery, one needs to maintain a clear distinction among the characteristics
The Geochemistry of Mass Extinction
of the global biota: their biodiversity (affected by differences in origination and extinction rates) and ecosystem function (guild structure, complexity of interactions, and productivity). Geochemical records reflect attributes of ecosystem function, not biodiversity; low-diversity recovery faunas and floras may support preevent productivities. Thus, geochemical and biodiversity recovery intervals are interdependent but not equivalent and may not be of equal duration. From the biologic point of view, there is an inevitable lag between peak extinction rates and peak origination rates, and the durations and underlying causes of the lags are topics of debate. Both intrinsic (e.g., the fact that ecospace is created as biodiversity increases, producing positive feedback) and external (environmental) constraints are possible. Kirchner and Weil (2000) performed a time series analysis of extinction- and originationrate data and concluded that the lag is approximately 10 My and independent of the magnitude of the event. Erwin (2001) raised the possibility that the 10 My lag may be an artifact of the coarseness of the timescales utilized and discussed possible environmental and ecologic limits on rate of recovery from mass extinction. The improved Paleobiology Database (Alroy et al., 2008) indicates that extinction rates decline markedly and origination rates increase immediately after mass extinctions but that >10 My lags in diversity recovery have occurred after the largest of the mass extinctions (e.g., the end-Permian) (Alroy, 2008). The comparison of the geochemical records of the five major mass extinctions of the Phanerozoic reveals few commonalities. Most, but not all, exhibit sharp drops in the carbon isotopic composition (d13C) of (at least) the surface ocean, indicating substantial disruptions to the global carbon cycle. The end-P and F–F events are associated with indicators of widespread anoxia and enhanced pyrite burial (positive d34S excursions), whereas the Late Ordovician extinction occurred during a brief interlude of oxic conditions from general anoxia. Some are associated with sea-level transgressions from previous lowstands (end-P, end-T, and K–P), but the Late Ordovician and F–F occurred during sea-level falls. Long-term climates change across all events but span major coolings (Late Ordovician and F–F) to prominent warmings (end-P, end-T, and K–P). Evidence for extraterrestrial influence is strong for the K–P, weak for the end-T and Late Permian, and missing for the F–F and Late Ordovician. What these times have in common is that all were times of biotic and environmental change. Long-term trends toward extreme environmental conditions presaged the Late Ordovician, F–F, and end-P events, whereas the end-T and K–P seem to have been abrupt shocks to Earth system, perhaps belying an extraterrestrial cause, although both are associated with large igneous provinces (as is the end-P). However, even for the K–P extinction, there is indication of environmental and biotic change before the known impact event and mass extinction (e.g., Abramovich and Keller, 2002; Barrera, 1994; Keller et al., 1993).
6.12.2 Isotopic Records of the Major Mass Extinctions 126.96.36.199
Carbon Isotopic Record
The carbon isotopic composition of marine limestones is generally interpreted to reflect the rate of organic carbon burial
(Forg, mol year1), although a number of other factors are actually involved and may be as significant for particular events or trends. The time rate of change of the d13C of the ocean, do, can be expressed as (e.g., Kump and Arthur, 1999; see Chapter 9.16). ddo Fin ðdin do Þ Forg D ¼ dt M
where Fin is the combined inorganic carbon input from weathering and volcanism (mol year1), din is its isotopic composition, D is the isotopic difference between contemporaneous sedimented organic carbon (dorg) and carbonate carbon (dcarb, do), approximately 25%, and M is the combined oceanic and atmospheric inorganic carbon reservoir size (mol). Because both (dindo) and D are negative numbers, an increase in do with time indicates either a reduction in the rate of carbon input, an increase in the isotopic composition of the input, an increase in organic carbon burial, a decrease in D (D becomes more negative), or some combination of these. The burial of carbonate carbon does not enter into this equation explicitly, because its isotopic composition is close to, and usually simplified to be equivalent to, do. When considering values representative of long (My) intervals of Earth history, steady state can be assumed, and do can then be expressed as do ¼ din
Forg D Fin
Because at steady state, Fin equals the sum of Forg and the carbonate burial rate (Fcarb), Forg/Fin reflects the fraction of the burial flux, which is organic (forg). From this equation, then, one sees that an interpretation of elevated do as high rates of Forg assumes that Fin, D, and din are all unchanging in time. As argued in the succeeding text, during particular intervals of Earth history, that assumption may be seriously in error. With these caveats in mind, the overall Phanerozoic pattern in do suggests that forg generally increased through the Paleozoic with peaks in the Silurian and Carboniferous, declined through the Late Paleozoic through Jurassic, and then oscillated without net trend through the Mesozoic–Cenozoic (Figure 1(a); Veizer et al., 1999). The Late Ordovician and Late Devonian were times of stepwise increases in do and perhaps forg, whereas the end-T and K–P events occurred at times when do was near its Mesozoic–Cenozoic mean. Higherresolution studies of the extinction intervals reveal positive excursions during the Late Ordovician and F–F extinctions (Saltzman, 2005) and negative excursions at the end-P (e.g., Luo et al., 2011), end-T (Ward et al., 2001), and K–P (D’Hondt et al., 1998) boundaries.
Oxygen Isotope Record
The most complete oxygen isotopic record for the Phanerozoic is that of Veizer and colleagues (1999). Although there is considerable controversy regarding the extent to which the data have been compromised by diagenesis (e.g., Mii et al., 1997), the general correspondence between other climatic indicators and d18O is encouraging (Veizer et al., 2000; Figure 1(a); see Chapters 8.17 and 8.20). Toward that end, it is interesting to note that, with the exception of the Late
The Geochemistry of Mass Extinction
8 P TT J
P TT J
0.7085 25 0.708 20
15 10 600 (b)
d 34S (CDT)
d 18O (PDB)
d 13C (PDB)
0.707 0.7065 500
Figure 1 Phanerozoic isotopic records. (a) C and O and (b) S and Sr (Adapted from Veizer J, Ala D, Azmy K, et al. (1999) 87Sr/86Sr, d13C and d18O evolution of Phanerozoic seawater. Chemical Geology 161: 59–89, (1999; see also http://mysite.science.uottawa.ca/jveizer/isotope_data/ index.html); Strauss H (1999) Geological evolution from isotope proxy signals – sulfur. Chemical Geology 161: 89–101). Vertical lines show the timing of extinction events.
Ordovician, four of the big five Phanerozoic mass extinctions occurred during times with relatively depleted d18O values (warm intervals). In detail, one finds that the Late Ordovician was a local peak in d18O, presumably reflecting the cool conditions of the time and, for the Hirnantian, the growth of continental ice sheets. The Late Devonian ushered in a cooling trend. A similar trend in d18O was initiated in the Late Permian and continued through the Triassic and Jurassic, but interpreting this in terms of gradual cooling is inconsistent with other climate indicators (e.g., Frakes et al., 1992).
Sulfur Isotopic Record
The basis of interpretation of the sulfur isotopic record is essentially identical to that for carbon; pyrite sulfur simply substitutes for organic carbon in the burial term of eqn . However, the long residence time of sulfur (sulfate) in the ocean ( 50 My; Holser et al., 1988) means that the time rate of change term on the left side of eqn  is important and cannot be neglected and a steady state cannot be assumed (i.e., eqn  cannot be used for the S cycle). The early Paleozoic exhibits heavy d34S values, and d34S declines more than 20% through the Paleozoic reaching a minimum of about 10% just before the end-P boundary (Figure 1(b)). A sharp rise in d34S in the earliest Triassic foreshadows the general trend through the rest of the Phanerozoic,
with d34S rising by about 10%. Apparently, pyrite sulfur burial rates were high in the early Paleozoic but declined through the Paleozoic, perhaps in response to an increasing proportion of organic carbon burial on land. A general decline in terrestrial coal basins from the Carboniferous to recent may be the explanation for the increase in d34S from the Late Paleozoic to now. Mass extinctions occurred during times of both heavy (Late Ordovician and Late Devonian) and light (end-P, end-T, and K–P) d34S. Large positive excursions in d34S followed the endP, end-T, and perhaps the Late Ordovician (Goodfellow et al., 1992) events, while a dramatic decline in d34S was initiated in the Late Devonian. The K–P event had no notable effect on the sulfur isotopic composition of the ocean. To the extent that d34S and d13C can be interpreted in terms of changes in the burial proportions of pyrite sulfur and organic carbon, consideration of both records together may shed some light on causal factors and/or environmental responses to these extinction events. The Late Ordovician and Late Devonian events were followed by positive excursions in d13C and either no or negative excursions in d34S. A positive excursion in d13C should be accompanied by a positive excursion in d34S if enhanced burial of marine organic carbon was the cause because C and S contents of marine sediments tend to covary (Berner and Raiswell, 1983). The S isotope records, instead, suggest that terrestrial organic carbon burial was enhanced during the Silurian and Carboniferous (Kump, 1992). That the Carboniferous was a period of enhanced terrestrial organic carbon burial is well established in the literature (e.g., Berner and Canfield, 1989). Perhaps the initial establishment of terrestrial ecosystems in the Silurian was responsible for the isotope signature of terrestrial burial during this period. In contrast, the positive d34S excursions after the end-P and end-T extinction events are best interpreted in terms of enhanced marine pyrite sulfur (and organic carbon) burial (e.g., Isozaki, 1997) but reduced global organic carbon burial. These excursions significantly postdate the extinction events, but because of the long response time of d34S, one should focus on the slope of its curve rather than the absolute value to get a sense of fluxes (e.g., Richter and Turekian, 1993).
Strontium Isotope Record
The strontium isotopic record of seawater reflects the relative magnitude of radiogenic (crystalline-rock weathering) and nonradiogenic (basaltic weathering and high- and lowtemperature seafloor reactions) sources to the ocean (e.g., Palmer and Elderfield, 1985). Periods of low 87Sr/86Sr are generally interpreted to reflect periods when seafloor hydrothermal activity was high, whereas periods of high ratios are generally interpreted to reflect times of elevated rates of continental weathering. In reality, the Sr cycle has multiple influences that are poorly constrained by the isotopic record alone (e.g., Kump, 1989). For the Sr isotope record, the present is the key to the Cambrian; in between, the 87Sr/86Sr ratio is seen to fall in a roller-coaster fashion through the Paleozoic and Early Mesozoic to a minimum in the Late Jurassic and then to rise in steps to the present (Figure 1(b)). The Late Ordovician and end-P events occurred at times of rapid increase following minima in
The Geochemistry of Mass Extinction
Sr/86Sr and thus marked times of significant change in the Sr cycle. In contrast, the Devonian and end-T events occur during local maxima in the ratio. The K–P event occurred contemporaneously with a local maximum in the ratio, on both the multimillion and 105-year timescales (MacLeod et al., 2001).
6.12.3 Interpreting the Geochemical Records of Mass Extinction 188.8.131.52
The first of the mass extinctions of the Phanerozoic occurred during the last stage of the Ordovician period, the Hirnantian, at c.440 Ma. The extinction appears to have occurred in two pulses (see Sheehan, 2001 for a thorough review of this event). The first was associated with a glacio-eustatic sea-level fall of 70–100 m as ice sheets developed on Gondwana, which at the time was situated at the South Pole. Large expanses of tropical to subtropical epicontinental seas replete with carbonate platforms, and diverse benthic and planktonic faunas became subaerially exposed. Marginal anoxic zones became fully oxygenated as oceanic mixing rates apparently intensified. After this first wave of extinction, a more cosmopolitan ‘Hirnantian’ fauna evolved, only to suffer considerable losses at the end of the glaciation as sea level rose, and shelfal anoxia was reestablished. The marine carbon isotope record provides some clues to the environmental causes and consequences of this event, but diverse interpretations have been published (e.g., Brenchley et al., 1994, 2003; Kump et al., 1999; Wilde et al., 1990; Young et al., 2009). Isotopic data from around the world (Young et al., 2009) confirm that a global positive excursion of 5–7% brackets the Hirnantian glaciation and extinction events (Brenchley et al., 1994; Long, 1993; Middleton et al., 1991; Wang et al., 1997). The glaciation apparently was confined largely to the Hirnantian and thus of short duration (0.5 My; Brenchley et al., 1994). In contrast, other Phanerozoic extinctions are associated with negative excursions and have generally been interpreted to represent the loss of surface-water productivity (photosynthesis discriminates in favor of 12C, creating 13C-enriched surface waters; see following text). Brenchley et al. (1994, 2003) proposed that the positive d13C excursion reflected increased marine biologic productivity promoted by more vigorous mixing of the ocean during glaciation. The resultant drawdown of atmospheric CO2 furthered the glaciation through positive climatic feedback. Alternatively, CO2 drawdown occurred before the glaciation in response to Taconic orogeny and associated enhanced weatherability of the continents (Kump et al., 1999; Young et al., 2009). A climatic threshold was reached, which allowed for the establishment and autocatalytic growth of Gondwanan ice sheets through positive ice–albedo feedback. The spread of ice reduced the exposure area and thus the weathering rate of silicate rocks (the long-term sink for atmospheric CO2), so CO2 levels rose. The limestonedominated rivers provided isotopically heavy carbon to the oceans (compared to preexcursion rivers, which had a higher proportion of carbon derived from shale-derived fossil-carbon weathering at higher latitudes), driving the positive excursion and perhaps masking the effects of the loss of surface-water
productivity on d13C. Ultimately, high CO2 created a sufficiently strong greenhouse effect to overcome the cooling effects of high albedo, and the ice sheets collapsed. As Sheehan (2001) points out, the two hypotheses are not mutually exclusive, although they present contrasting predictions for the time course of atmospheric CO2 through the event. Weak proxy evidence of atmospheric CO2 rise through the event, in the form of a reduction in the isotopic difference between limestone and kerogen d13C, is supportive of the weathering hypothesis (Kump et al., 1999; Young et al., 2009). On the other hand, the apparent simultaneous increase in d13C and d18O of carbonates at the onset of the event has been argued by Brenchley et al. (2003) to be inconsistent with the weathering hypothesis in that according to that model, the oxygen isotope shift (ice-sheet growth and cooling) should begin before the carbon isotope excursion. However, as Brenchley et al. (2003) acknowledge, the predicted lag could be shorter than one’s ability to resolve it. Lack of a deepwater (benthic) carbon isotope record hinders an assessment of the extent to which pelagic ecosystem function (biologic pumping of organic matter, nutrients, and trace metals to the deep sea) was disrupted during the event. In contrast, as discussed in the succeeding text, there is clear evidence in the form of a collapsed d13C gradient from surface to deep sea for a shutdown of the biologic pump during the K–P event. The fossil record indicates that, although the pelagic biota was certainly not immune to Late Ordovician extinction, recovery of abundance and diversity was rapid (Sheehan, 2001). Fossil soil carbonate d13C (e.g., Cerling, 1991; Mora et al., 1991) and/or the concentration and d13C of carbonate substituted in goethites in ancient soils (Yapp and Poths, 1992) have been proposed as pCO2 barometers. In the case of the Late Ordovician (for which there are very few paleosols), Yapp and Poths (1992) derived a very high estimate of paleopCO2 ( 16 times present atmospheric level). Presuming that the paleosol was deposited at the time of glaciation, these high CO2 levels seem paradoxical (Kasting, 1992). Climate models show that glaciation can occur at high pCO2 (Crowley and Baum, 1991) under certain paleogeographical conditions (e.g., a large continent with its coastline at a pole), but Gibbs et al. (1997) argued for a maximum pCO2 of about 8–10 present under Ordovician conditions of paleogeography and (reduced) solar luminosity (see also Chapter 6.11). If the paleosols record CO2 levels at the end of the glacial period, the paradox of high atmospheric pCO2 during glaciation is reconciled; under the weathering hypothesis, pCO2 rises during the glaciation to levels sufficient to overcome the ice–albedo effect. Further support for the weathering hypothesis for Late Ordovician glaciation comes from the Sr isotopic record. A multimillion-year decline in the marine record of 87Sr/86Sr in limestones is reversed in the early Late Ordovician (Veizer et al., 1999), a likely consequence of collisional tectonics that initiated at this time and continued (as did the rise in 87Sr/86Sr) through the Silurian (Richter et al., 1992). The increase (from 0.7078 to 0.7088) rivals the Cenozoic rise, which has been attributed to the Himalayan uplift and associated with the progressive cooling leading to Quaternary glaciation (Raymo et al., 1988). The large shifts in Sr isotopic composition of the ocean probably reflect a substantial increase in the supply of
The Geochemistry of Mass Extinction
radiogenic Sr facilitated by the unroofing of older continental crust during orogeny (Richter et al., 1992). Climatic cooling compensated for tectonically increased weatherability of the continents via feedbacks between climate, CO2, and global weathering (Kump and Arthur, 1997). Recent work with clumped oxygen isotopes reveals an enigmatic temperature and ice volume history that is challenging to explain in the context of Pleistocene glaciation and climate simulations (Finnegan et al., 2011). The data suggest that, although ice-sheet growth began in the Katian Stage of the Late Ordovician, tropical temperatures remained exceedingly warm (by modern standards, 32–37 C) until the earliest Hirnantian, when they fell by 5 C and then rebounded rapidly, before the end of the Hirnantian, and well before the total collapse of the Hirnantian ice sheets. Presuming these climate changes were driven by CO2 fluctuations, they tend to favor the productivity hypothesis for the carbon isotope excursion over the weathering hypothesis. A parallel positive shift in d34S of pyrites (Goodfellow et al., 1992; Zhang et al., 2009) occurs during the Late Ordovician and is superimposed on a longer-term Ordovician–Silurian increase in d34S of the ocean. Generally, one interprets parallel shifts in d13C and d34S as indicative of a marine organic carbon burial event (Gill et al., 2011; Kurtz et al, 2003), perhaps driven by enhanced marine productivity and/or anoxia development in the deep ocean. Efforts to find trace-metal evidence of extraterrestrial impact at the Ordovician–Silurian boundary have been unsuccessful (e.g., Orth et al., 1986; Wang et al., 1992). Peaks in Ir abundance at the boundary have been linked to reductions in sedimentation rate; the persistent cosmic source of Ir is otherwise diluted by high terrigenous or biogenic sedimentation. Overall, the Late Ordovician extinction appears to be the result of purely terrestrial phenomena. High sea-level stands of the Early Paleozoic allowed for animal diversification on shallow-water, epicontinental carbonate platforms that proved, however, to be highly sensitive to glacio-eustatic effects on ecospace availability and lateral shifts in the oxic/anoxic interface. Tectonic activity facilitated the establishment of Gondwanan ice sheets, which robbed the shallow seas of water, leading to extinction, establishment of recovery ecosystems, and then the destruction of these as the ice sheets melted, perhaps catastrophically.
The Late Devonian was a time of widespread, shallow epicontinental seas that supported abundant and diverse warmwater metazoan communities. This biotic bliss was terminated in a series of extinctions extending over perhaps 3 My (McGhee, 1996), which had lasting effects on reef-building stromatoporoids, corals, brachiopods, and fish (e.g., McLaren, 1982). Reefs recovered in the Famennian but, at least in Western Australia, were dominated by cyanobacteria (Playford et al., 1984). Global cooling may have played a role in the preferential extinction of warmwater faunas (including coral reefs; McGhee, 1996). Widespread anoxia has also been invoked to explain this interval of elevated extinction (e.g., Joachimski and Buggisch, 1993). Associated eutrophication is invoked by Murphy et al. (2000) to explain the demise of Devonian
carbonate platforms. Their hypothesis is supported by substantial increases in the C/N and C/P ratios of organic matter preserved in the Kellwasser deposits. Preferential release of nutrients is argued to occur during early diagenesis, when overlying waters are anoxic, supporting high levels of productivity in the water column (Van Cappellen and Ingall, 1994, 1996). Anoxia was certainly a characteristic of the Late Devonian, but in the Western United States, based on geochemical proxies, anoxia ended 6 m below (100 ky before) the major F–F extinction. Bratton et al. (1999) discuss the possibility that this was a local phenomenon and that anoxia persisted through the F–F boundary elsewhere but favor the alternative hypothesis that other sections suffered depositional hiatus or erosion of the latest Frasnian sediments. A positive d13C excursion (in both carbonate and organic carbon) began in the Frasnian but continued well into the Famennian (Wang et al., 1996). A positive pyrite-S isotope excursion also occurred at this time. If these excursions indicate enhanced organic carbon and pyrite sulfur burial under widespread anoxic conditions, then it would seem that such conditions persisted well beyond the F–F boundary extinction. Wang et al. (1996) identified a brief negative d13C excursion at the F–F boundary in Alberta, Canada, which may reflect the temporary loss of the biologic pump. They argue that productivity collapse was the result of an asteroid/comet impact at the F–F boundary, a time otherwise under biotic stress as the result of widespread warmth and anoxia. Murphy et al. (2000) detail a similar carbon isotope stratigraphy, in this case, in organic matter, with two peaks from a baseline of d13C ¼ 31 to 27%, representing the Lower Kellwasser (Frasnian) and Upper Kellwasser (F–F boundary) episodes of black-shale deposition in Europe (Figure 1(c)). Between these two events, d13C drops to 32% just prior to the F–F boundary. Perhaps, this is the brief negative excursion of Wang et al. (1996). Correlation between organic carbon and carbonate carbon d13C records is imprecise because of the additional possibility of productivity or atmospheric pCO2-induced variations in isotopic fractionation that could generate phase offsets between the two records (Kump and Arthur, 1999). Nevertheless, in this case, based on paired analyses from the same section, the inorganic and organic records appear to be in phase, suggesting that atmospheric CO2 levels were sufficiently high that any changes did not significantly affect the isotopic fractionation that occurs during photosynthesis (Joachimski et al., 2002). The impact hypothesis for this extinction dates back to McLaren (1970). There is some support for extraterrestrial impact at this time. Iridium anomalies have been identified at or near the F–F (Playford et al., 1984; Wang et al., 1991) and Devonian–Carboniferous (Wang et al., 1993) boundaries. However, the Devonian–Carboniferous anomaly has no supporting evidence for impact, and elemental ratios are not chondritic. Redox changes have been invoked to explain this anomaly (Wang et al., 1993). Stronger evidence for impact is present for the F–F boundary (as summarized in Wang et al., 1996), including microtektites, meteoritic elemental ratios, known impact craters, and high-energy (tsunami) deposits. However, reported Ir enrichments at the stratotype area for the F–F boundary in Southern France were not subsequently substantiated (Girard et al., 1997), and no F–F Ir anomaly was found in New York State (McGhee et al., 1984).
The Geochemistry of Mass Extinction
The current paradigm focuses on transgression at the F–F boundary and the spread of anoxia across epicontinental seas to explain the extinction (e.g., Bond and Wignall, 2009). In Bond and Wignall (2009) analysis, cooling and sea-level fall preceded the event. The selectivity of the event, in particular the fact that sessile calcifiers, such as the reef-building corals, were particularly hard-hit, is consistent with a mechanism that provides metabolic stress through low oxygen and, presumably, elevated CO2 concentrations (cf. the Knoll et al., 1996 explanation for the end-Permian event).
The largest extinction event of the Phanerozoic occurred in the latest Permian, a time when both shallow and deep marine environments appear to have experienced widespread anoxia. As a result, anoxia has figured prominently into proposed extinction mechanisms for this time, although models for extinction that invoke multiple causalities are currently in favor (e.g., Erwin, 1993, 1995, 2006; Kozur, 1998). The latest Permian isotopic record displays an abrupt negative excursion in carbonate and organic carbon isotopes (recently reviewed by Korte and Kozur, 2010; see also Luo et al., 2010) and a substantial increase in the sulfur isotopic composition of the ocean as recorded in evaporite sulfate minerals and trace sulfate in carbonate minerals in the early Triassic following a minimum value in the Late Permian (e.g., Kajiwara et al., 1994; Scholle, 1995; Strauss, 1997, 1999; Riccardi et al., 2006; recent summary by Luo et al., 2010). These proxies may represent the widespread development of oceanic anoxia and the establishment of strong chemical stratification of the ocean (Gruszczynski et al., 1989). An unusual chemistry (e.g., anoxia and widespread carbonate supersaturation) extended into the early Triassic and may have been a contributor to the long recovery interval (e.g., Hallam, 1991; Woods et al., 1999). The coincidence of Siberian flood basalt emplacement (Reichow et al., 2009; Renne et al., 1995) with the anoxia and extinction event suggests a causative role for volcanism (see also Kozur, 1998; Svensen et al., 2009). Tentative evidence for a cometary impact exists in the form of noble gases in fullerenes (Becker et al., 2001), unusual Ni-rich grains (Kaiho et al., 2001), purported chondritic meteorite fragments in Late Permian sediments (Basu et al., 2003), and impact craters whose broad ranges of allowable ages include the Late Permian (Woodleigh structure: Mory et al., 2000a; Bedout structure: Becker et al., 2004; and Wilkes Land structure: von Frese et al., 2009). However, the fullerene and other extraterrestrial results (Becker and Poreda, 2001; Farley and Mukhopadhyay, 2001; Isozaki, 2001; Koeberl et al., 2004) and, moreover, the impact crater interpretation of the Woodleigh (cf. Mory et al., 2000b; Reimold and Koeberl, 2000) and Bedout have been challenged. Numerical modeling indicates that Permian deepwater anoxia required either low atmospheric pO2 or warm bottomwater source regions, together with elevated oceanic nutrient (phosphate) concentrations (Hotinski et al., 2001; Meyer et al., 2008). Warm source regions acquire lower oxygen concentrations at equilibrium with the atmosphere before sinking, providing less oxygen to deep waters. When upwelled to the
surface, higher phosphate concentrations in deep waters intensify the biologic pump and thus increase O2 demand in deep waters (see Chapter 8.4). These factors prove to be much more important than the sluggish circulation itself, which not only does reduce O2 supply to deep waters but also reduces the strength of biologic pumping and thus the O2 demand on deep waters (e.g., Hotinski et al., 2000, 2001; Meyer and Kump, 2008; Ozaki et al., 2011; Sarmiento et al., 1988). In fact, another model failed to generate anoxia under reasonable Permian conditions (Zhang et al., 2001). Subsequent analysis of the Zhang et al. (2001) results (Hotinski et al., 2002) indicated that surface forcings in their model generated a more vigorous circulation of cooler water than in the model of Hotinski et al. (2001; cf. Winguth and Maier-Reimer, 2005). The transgression of anoxic deep waters with sea-level rise may have been the direct kill mechanism (Cirilli et al., 1998; Hallam, 1991; Wignall, 1990; Wignall and Hallam, 1992, 1993; Wignall and Twitchett, 1996). Alternatively, anoxic deep waters may have undergone periodic catastrophic upwelling, induced by cooling or other surface forcings, causing significant transients in CO2 and perhaps H2S surface-water concentrations (Knoll et al., 1996). These concentrations would have induced CO2 toxicity (hypercapnia), especially in calcifying organisms, and could have been lethal. Selectivity of marine extinctions for those organisms especially sensitive to high aqueous CO2 concentration supports the hypercapnia explanation for the extinction (Knoll et al., 1996). Terrestrial extinctions could have been produced by the climate changes attendant upon the equilibration of the atmosphere with surface waters enriched in CO2. This notion is supported by the evidence that widespread deepwater anoxia developed considerably before the main extinction event (Isozaki, 1997). However, a physical mechanism, such as enhanced upwelling, may not have been necessary. Kump et al. (2005) show that a threshold exists during the buildup of H2S in deep waters (due to eutrophication) whereby the upwelling flux of H2S comes to exceed the influx of O2 from the atmosphere, and the surface layer becomes euxinic. Abrupt shifts into euxinia would be focused in regions of strong physical upwelling because the threshold H2S concentration would be lower. Such ‘chemocline upward excursion’ events have been documented in the Late Permian (Riccardi et al., 2006, 2007) and reproduced using Earth system models of intermediate complexity (Meyer et al., 2008). An additional implication of CUE events is the release of H2S to the atmosphere. The initial 1-D atmospheric modeling of Kump et al. (2005) indicated that the expected flux of H2S from widespread photic-zone euxinia could deplete the hydroxyl content of the troposphere, leading to toxic levels of H2S and abrupt increases in methane concentrations (with the loss of the hydroxyl sink) and destruction of the ozone layer. Kump et al. (2005) argued that poisoning by H2S would be a satisfying way to link the marine and terrestrial extinctions, presuming that the CO2-induced warming would likely be insufficient. Evidence for genetically deformed fossil spores (Visscher et al. 2004) seemed to reinforce this hypothesis, as did the growing evidence from the presence of the green-sulfur bacterial biomarker isorenieratane in end-Permian sediments for widespread photic-zone euxinia (Grice et al., 2005; Hays
The Geochemistry of Mass Extinction
et al., 2007). However, more sophisticated 2-D (Beerling, 2002) and 3-D (Lamarque et al., 2006) modeling indicates that the catastrophic implications of the Kump et al. (2005) simulations likely overestimated the buildup of H2S for the range of air–sea fluxes evaluated and that much less toxic levels likely resulted from CUE events. These results do not rule out the possibility of patches of H2S-rich air developing episodically in coastal zones and wafting across the continents, especially during hurricanes, which bring up subsurface waters. Hypercanes (especially large and intense hurricanes expected to occur during extreme warm intervals of Earth history; Emanuel et al., 1995) may have occurred during the greenhouse warmth of the Late Permian (J. Kiehl, pers. comm.). Halides injected into the atmosphere from explosive Siberian Trap volcanism penetrating sedimentary deposits, including evaporites, may have been effective in destroying the ozone layer (Visscher et al., 2004). Research on the climatic and carbon isotopic effect of destabilization of methane clathrates (e.g., Dickens et al., 1995, 1997, 2001, 2011) during a (possibly) similar event at the Paleocene–Eocene boundary suggests yet another hypothesis for the end-Permian disaster: Siberian Trap volcanism led to warming and rearrangement of ocean circulation patterns, bringing warm intermediate waters into contact with continental shelf sediments, leading to the catastrophic release of methane to the ocean, and generating anoxia, negative d13C excursions (de Wit et al., 2002; Krull and Retallack, 2000), and enhanced global warming (Wignall, 2001). Numerical modeling of the end-Permian carbon cycle supports the hypothesis of methane-driven isotope shifts (Berner, 2002) for some, but not all, of the fluctuations during this interval (Payne and Kump, 2007).
were constant over this interval (cf. Beerling, 2002; Tanner et al., 2001, 2002). The geochemical record of the end-T extinction is rather limited, but recent additions have been made that are producing a clearer picture of the event. Carbon isotope records from organic matter in boundary sections from Canada (Ward et al., 2001), England, and Greenland (Hesselbo et al., 2002) and carbonates in Hungary (Palfy et al., 2001) document a negative excursion of 2–4% centered on the boundary. Ward et al. (2001) ascribed this to a collapse of marine productivity and likened the response to the end-P and K–P extinction records. There is no apparent shift in the Sr isotopic composition of the ocean across the event (Hallam, 1994). A modest Ir anomaly, together with a substantial increase in fern spore abundance, occurs precisely at the end-T boundary (Olsen et al., 2002), indicating that a meteorite impact may have triggered the mass extinction. However, volcanism could also explain the Ir anomaly (Whiteside et al., 2010). In the most comprehensive study to date, Whiteside et al. (2010) make a compelling case that CAMP volcanism triggered the extinction. The negative carbon isotope excursion is apparent and synchronous in both terrestrial and marine organic matter and with the initiation of CAMP volcanism in Morocco. Stomatal density in plant fossils from the same sections indicates that atmospheric pCO2 increased during the event. All of these results are consistent with volcanism as the trigger for the end-T extinction. However, the kill mechanism remains unknown. Anoxia does not seem to be the cause (Wignall and Bond, 2008); perhaps, halogens injected into the stratosphere by CAMP volcanism were the culprit, drawing on the Beerling (2002) explanation for the end-Permian mass extinction.
Of the Phanerozoic ‘big five,’ the mass extinction that occurred at the Triassic–Jurassic boundary has received the least attention from geoscientists until fairly recently (e.g., Deenen et al., 2010; Mander et al., 2010; McElwain et al., 2009; van de Schootbrugge et al., 2009; Whiteside et al., 2010). The event affected the diversity of both terrestrial and marine ecosystems and may have created the opportunity for the rise to dominance of dinosaurs by selective extinction of the nondinosauran competitors (Olsen et al., 2002). Its timing coincides with the emplacement of a large igneous province (the Central American Magmatic Province) and associated volcanic activity (Hesselbo et al., 2002; Marzoli et al., 1999). Interestingly, the event appears to have initiated on land some several hundred thousand years before it did in the ocean, suggesting a trigger to which terrestrial ecosystems were most sensitive, followed by prolonged environmental change that adversely affected marine ecosystems (Palfy et al., 2000). Others have challenged this interpretation, arguing that the events on land and sea were synchronous (Hesselbo et al., 2002). Paleobotanical data (stomatal density) suggest that atmospheric CO2 levels increased fourfold across the boundary, perhaps increasing leaf temperatures to lethal limits (McElwain et al., 1999), but soil carbonate isotope proxies suggest that CO2 levels
The fossil record shows that the species composition of terrestrial and marine ecosystems suffered a nearly complete turnover at the Cretaceous–Paleogene boundary. Marked iridium anomalies (e.g., Alvarez et al., 1980; Kyte et al., 1985) and other indicators of asteroid or comet impact (several papers in Sharpton and Ward, 1990), including the Chicxulub impact crater (Hildebrand et al., 1991), have provided compelling evidence for an extraterrestrial cause of this extinction (Schulte et al., 2010). However, the carbon isotope record of this event has shed the most light on how this mass extinction affected the operation of the Earth system. Through preferential incorporation of 12C into organic matter during photosynthesis, the biologic pump establishes an isotopic gradient between the surface and deep of 1–3% (e.g., Kroopnick, 1974; see also Chapters 8.4 and 8.18). The isotopic record of the latest Cretaceous to earliest Paleogene, however, shows an essentially complete collapse of this gradient, which for some sites is interpreted to have lasted from hundreds of thousands (Zachos et al., 1989) to 3 Ma (D’Hondt et al., 1998). Hsu and McKenzie (1985) referred to this state as the ‘Strangelove Ocean’ because it was their conclusion that it represented a near cessation of primary production in the surface ocean. This interval is now recognized in other regions of the world (e.g., Barrera and Keller, 1990; D’Hondt et al., 1998; Keller and Lindinger, 1989; Stott and
The Geochemistry of Mass Extinction
Kennett, 1989) and among other groups of organisms, including mollusks, that lived in offshore shelf environments (Hansen et al., 1993). A number of suggestions for the cause of the d13C gradient collapse have been made, including cessation of biologic productivity or enhanced oceanic overturn (Zachos et al., 1989), but Kump (1991) has shown that the loss of the biologic pump is the explanation most consistent with mass balance constraints on the carbon isotopic system. Interestingly, it is conceivable that productivity in the surface ocean remained high, perhaps aided by an explosion in the abundance of stress-adapted plankton (Hollander et al., 1993; Percival and Fischer, 1977). If so, the loss of the biologic pump would have to be explained by a loss in the ability of the ocean system to aggregate finegrained organic matter into large particles or provide the ballast (usually dense, biogenic mineral material, especially CaCO3) to facilitate sinking (Armstrong et al., 2002). Any zooplankton that appeared soon after the event may not have created sufficiently large fecal pellets to facilitate sinking (D’Hondt et al., 1998). A dearth of coarser-grained calcareous material in their fecal pellets may have also contributed to a weak biologic pump. There are other geochemical tracers of biosphere response to mass extinction besides carbon isotopes. For example, tracemetal contents of K–P sediments may provide additional evidence about the extinction itself and the ensuing ‘Strangelove Ocean’ recovery interval (Hsu et al., 1982; Officer and Drake, 1985; Vogt, 1972). Erickson and Dickson (1987) calculated that the metal burden of a 10 km meteorite would increase the metal content of the ocean several fold. They used the presentday residence times of the elements to argue for the rather fast removal of the short residence-time elements (e.g., Fe) and longer times, thousands of years, for others (e.g., Ni and Cu). However, the major removal mechanism for many metals is in association with the biologic pump. Thus, the enrichment of metals as a direct result of meteorite vaporization and solubilization (and perhaps as the result of an interval of intense continental weathering due to nitric acid rain; Macdougall, 1988) is likely to persist much longer in the absence of the pump. Over even longer intervals of time, the metal content of the oceans would continue to rise as rivers, aerosols, and hydrothermal activity provided metals to the ocean at a rate much faster than their rate of removal in the absence of the biologic pump. Many short-lived elements today might have become long-lived during the ‘Strangelove Ocean’ interval, with ocean mixing homogenizing their distributions. Thus, the persistence of the ‘Strangelove Ocean’ in the face of high rates of diversification and expansion of stress-related biota may have been due to the persistence of toxic concentrations of metals (Leary and Rampino, 1990). Jiang et al. (2010) invoked the Erickson and Dickson (1987) hypothesis to explain the timing and spatial distribution of extinction intensity among calcareous nannoplankton during the K–P event. Lower extinction intensity and more rapid recovery of diversity in the Southern Hemisphere relative to the North was attributed to a northward oblique impact shedding ejecta into the northern oceans, where the associated metals were solubilized and preferentially retained due to the likely pattern of meridional overturning of the K–P ocean. Based on estimated extents
of solubilization and residence times, Fe, Al, Co, Mn, Ni, Cr, Cu, Pb, Hg, Zn, Ag, and Cd would all be strongly enriched in surface waters initially; after a century, Fe, Al, Co, Mn, Ni, and Cr would remain more than 10 enriched relative to today, and even after 10 000 years, Cr and Cu would still be enriched by factors of 5–10-fold, leading to the suggestion that the event be called the ‘Erin Brockovich’ rather than the ‘Strangelove Ocean.’ Direct proxies of water-column metal contents have only been established in one case for the K–P boundary (Stott and Delaney, 1988). These investigators found a positive excursion in the Cd/Ca ratio in benthic foraminifera at site 690C (Weddell Sea, Antarctica) but just before the K–T boundary. There was no apparent change in the ratio across the boundary, which Stott and Delaney (1988) interpreted to indicate that there was little change in productivity at this site as a result of the extinction.
Summary with Extensions
It can safely be said that there are no universal geochemical precursors or responses to extinction events in the Phanerozoic (Table 1). The Paleozoic extinctions are associated with indicators of anoxia; the end-Permian, end-Triassic, and K–P events coincide with significant volcanic events, and there is weaker evidence for volcanism in the Late Devonian (Wignall, 2001); only the K–P event has conclusive evidence for asteroid or comet impact. In considering the causes and consequences of mass extinctions, it is important to distinguish the trigger for the event from the kill mechanism; these can be two quite distinct phenomena. Knoll et al. (2007) provide a clear basis for separating the two: “A kill mechanism is the physiologically disruptive process that causes death, whereas a trigger mechanism is the critical disturbance that brings one or more kill mechanisms into play.” They go on to point out that the focus of much research has been on identifying triggers, and geochemistry certainly comes into play in this endeavor. Volcanic eruption emerges as a likely trigger for most of the Phanerozoic extinctions and may have served as a kill mechanism as well through its direct impact on stratospheric chemistry (Beerling, 2002). If warming is initiated by volcanic CO2-induced marine anoxia, hypercapnia (CO2 toxicity; Knoll et al., 1996) or H2S poisoning (Kump et al., 2005) may have resulted, and the latter can be recorded in the biomarker record of the event (e.g., Grice et al., 2005). Collapse of biologic productivity (the biologic pump) is a likely consequence of mass extinction, but this is recorded unambiguously as a reduction in the carbon isotope gradient in the ocean for only the K–P event. Negative d13C excursions during the end-P and end-T events may also be a reflection of biologic pump collapse, and evidence is emerging to support this contention (Meyer et al., 2011) despite a lack of a benthic fossil CaCO3 record for the pre-Jurassic. Convincing evidence that trace-metal poisoning from volcanism or asteroid impact has yet to appear, but that and persistent anoxia have been invoked to explain the delay in recovery after some extinctions. Resolution of these uncertainties will clearly require continued, coordinated stratigraphic sampling
The Geochemistry of Mass Extinction
Geochemical and environmental phenomena associated with mass extinction
87 Sr/ 86Sr excursion
Global temperature change
Cretaceous– Paleogene Triassic– Jurassic End-Permian
(þ) (short term) (0)
N (but Fe–Si–Ni grains and fullerenes) N (but Y above and below F–F boundary) N
Generally (þ) but () right at F–F boundarya,c (þ) (first wave) and () (second wave)
(þ) (from Paleozoic minimum) (þ) (from Devonian minimum) (þ) (from Ordovician minimum)
Y (but may have ended just before F–F)e Y, before and after glaciation
" After Maastr. lowstanda " After Late Triassic # " From Late Permian lowstand #
# then "
# then "
Late Devonian (F–F) Late Ordovician
Keller et al. (1993). Olsen et al. (2002). c Wang et al. (1996). d Copper (1998). e Bratton et al. (1999). b
and geochemical and isotopic analysis of the fossils and their sedimentary matrix, together with the development of new proxies and improved geochronological approaches (perhaps including orbital cyclostratigraphic approaches; Rampino et al., 2002).
Acknowledgments The author acknowledges support from the NASA Astrobiology Institute and the NSF Geobiology and Environmental Geochemistry and Continental Dynamics Programs.
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