- © 2005 by the Mineralogical Society of America
In the past 300 million years, there has been a near-perfect association between extinction events and the eruption of large igneous provinces, but proving the nature of the causal links is far from resolved. The associated environmental changes often include global warming and the development of widespread oxygen-poor conditions in the oceans. This implicates a role for volcanic CO2 emissions, but other perturbations of the global carbon cycle, such as release of methane from gas hydrate reservoirs or shut-down of photosynthesis in the oceans, are probably required to achieve severe greenhouse warming. The best links between extinction and eruption are seen in the interval from 300 to 150 Ma. With the exception of the Deccan Trap eruptions (65 Ma), the emplacement of younger volcanic provinces has been generally associated with significant environmental changes but little or no increase in extinction rates above background levels.⇓
Can large igneous province (LIP) eruptions cause environmental and climatic effects that are sufficiently severe to cause extinctions? The answer to this question has changed significantly over the past two decades, from “probably no” to “probably yes.” The change of view is due to advances in radiometric dating of both extinction events and the ages of LIPs ⇓(see Courtillot and Renne 2003). The close age correspondence often demonstrated by dating now provides the most compelling link between the two phenomena. Thus, four mass extinctions have occurred in the 300 million year interval between the Permian and the present day, together with a similar number of minor biotic crises. All of these crises coincide with LIP eruptions (Table 1). No other phenomenon shows such a 100% correlation; certainly not meteorite impact. However, not all LIP eruptions in this interval coincide with extinctions. For example, the especially large Paraná-Etendeka Province was erupted 133 Ma ago, in the early Cretaceous, at a time marked by extremely low extinction rates. Thus, the claim that all extinction events coincide with giant volcanic eruptions is well substantiated, but the converse that all such eruptions coincide with extinctions is not true.
TOWARDS A KILL MECHANISM
In general, LIP eruptions are associated with some or all of the following climatic and environmental effects:
Rapid global warming
Oceanic anoxia or increased oceanic fertilisation or both
A sharp decrease in the δ13C values recorded in limestones; this is usually interpreted as a record of methane release from gas hydrate reservoirs.
The Karoo-Ferrar eruptions around 180 million years ago present all of these features, and the Siberian Traps eruptions 250 million years ago show many of them. The other LIP eruptions show several features (Table 1). Volcanogenic cooling has also been proposed for several extinction events, but the evidence is insubstantial. The preponderance of evidence favours warming. This strongly suggests that CO2 emissions do all the damage, although in many extinction scenarios this effect is envisaged as a trigger (Fig. 1). Indeed other factors are almost mandatory given the volume of CO2 likely to be released during LIP eruptions (Self et al. this issue). The amount of CO2 released during the eruption of the largest LIPs is unlikely to have exceeded 1013 tonnes of CO2, with the amount released during individual flow events likely to have been at least two orders of magnitude lower (Wignall 2001). Thus, the gas released during a major flow of 1000 km3 (which may have occurred as frequently as every few thousand years) is unlikely to have greatly exceeded the current anthropogenic CO2 release rate of 25 × 109 tonnes per annum. This modern flux comes not even close to recreating the conditions during these ancient catastrophes. It is possible that LIP eruptions are associated with excessively CO2-rich volcanism, reflecting a mantle source still rich in volatiles. However, this has yet to be demonstrated. Alternatively, the volcanism may serve as a trigger for something else, such as the release of methane from clathrates buried at shallow depths beneath the seafloor. Methane is a much more effective greenhouse gas although it is rapidly oxidised to CO2 in the atmosphere. However, the tell-tale evidence for methane release - the rapid decrease of the 13C/12C ratio recorded in organic carbon and limestones, known as a negative δ13C anomaly (see glossary) - is seen less frequently than the other evidence for warming.
A Volcanic Greenhouse Scenario
Ideas concerning the role of volcanic gases have been developed primarily from events during the end-Permian and Early Jurassic mass extinctions (e.g. Pálfy and Smith 2000; Wignall 2001). It is proposed that the volcanic extinction mechanism is triggered by the release of CO2 during the eruption of the giant lava flows that form LIPs (Fig. 1). The resultant increase of atmospheric CO2 levels would have several deleterious consequences for the oceans. For example, an increase of CO2 concentrations in surface waters causes a pH decline and thus problems for carbonate-secreting organisms (Gattuso and Buddemeier 2000). This is known as a calcification crisis and is manifest as a decline in carbonate content in many sections. Global warming can also cause the development of oceanic anoxic events (and therefore marine extinctions). These are intervals of time when large areas of the oceans and shelf seas were either oxygen poor or oxygen free (anoxic). Modern oceans typically have around 5-6 ml of oxygen dissolved in a litre of water, but conditions become stressful for most organisms if values decline below 1.0 ml/L, and no metazoan life can survive below values of 0.3 ml/L. Low oxygen conditions are restricted to only a few small areas of modern oceans but they became much more widespread during oceanic anoxic events due to several feedback factors associated with global warming. First, warmer waters hold less dissolved oxygen than colder waters. Second, the ocean's circulation system is primarily driven by the temperature gradient between the equator and the poles, with deep circulation driven by the generation of cold and dense waters in the polar regions. This system slows down as polar waters warm up, thus decreasing the supply of oxygen to the ocean's deeper waters. A possible third factor may relate to the supply of nutrients from land, which will increase with global warming due to increased rainfall and runoff in a warmer, more humid climate. Increased nutrient flux to the seas will foster increased biological productivity, which in turn will decrease oxygen levels in sea water as the plankton biomass decays - the same phenomenon is seen in many modern shelf seas over-supplied with anthro-pogenic “nutrients” such as fertilisers and sewage. Evidence of increased global runoff during LIP eruptions is substantial and includes several lines of geochemical evidence, such as an increase in the trace metals rhenium and osmium (Ravizza and Peucker-Ehrenbrink 2003). However, many mass extinction events coincide with a collapse, not an increase, of primary productivity (Hallam and Wignall 1997), and this third factor may not be significant until after the mass extinction has run its course. Indeed its main significance may be as a vital negative feedback loop for drawing down atmospheric CO2 (Fig. 1). As already noted, there is debate as to whether volcanic CO2 emissions are sufficient on their own to cause these environmental changes; other phenomena such as gas hydrate release may also contribute to increasing greenhouse gas concentrations.
The volcanic greenhouse scenario is currently a “working hypothesis” for several marine extinction events, but its relevance to contemporaneous terrestrial extinction events has not been explored to any great extent. Severe global warming will obviously shift climatic belts and presumably restrict habitat area for the most cold-adapted communities. However, tropical habitats should benefit rather than suffer from such changes. Consequently most terrestrial extinction mechanisms focus on other aspects of LIP eruptions. Volcanic halogen emissions can potentially damage the ozone layer, thus raising the spectre of UV radiation as a contributory cause of extinctions. Mutant and deformed plant spores and pollen in the end-Permian extinction interval may be a sign of this radiation damage (Visscher et al. 2004); it is certainly evidence for extreme environmental stress, but it has yet to be established if such phenomena are a regular feature of LIP eruptions. Acid rain from volcanogenic sulphate aerosols is another potentially harmful effect of flood basalt eruptions, but there is, as yet, little direct evidence for this.
Volcanism-extinction scenarios have been developed primarily to explain end-Permian and Early Jurassic extinction events, and it is interesting to compare these with environmental changes seen during some other LIP eruptions of the past 300 Myr.
Central Atlantic Magmatic Province (CAMP) (200 Ma)
The realisation that the Central Atlantic Magmatic Province was both very large and of the right age to be implicated in the end-Triassic mass extinction suggested yet another important volcanism-extinction link (Marzoli et al. 1999). This extinction event has proved rather difficult to study, primarily because of a dearth of complete marine boundary sections. This paucity reflects the extremely low sea level at this time, which could of course have contributed to the marine extinctions (Hallam and Wignall 1999). Other changes at this time include a brief warming pulse, carbon isotope evidence for a significant release of methane and possibly a calcification crisis (Hautmann 2004). Marine anoxia was widespread, but this seems to have been the case both before and after the extinctions.
Clearly the CAMP volcanism may have been responsible for these environmental changes, but there are some problems associated with a CAMP-extinction link. First, there is the detailed timing. The sedimentary record of the Newark Basin in northeastern USA contains evidence for both terrestrial extinctions and flood basalt volcanism; however, the first basalt occurs somewhat above the extinction horizon (Wignall 2001). Furthermore, it is not at all clear that the end of the Triassic was marked by a sudden, single-pulse mass extinction (Hallam 2002). Extinction rates were high throughout the last few million years of the Triassic, suggesting a prolonged crisis that began considerably before the CAMP eruptions.
Paraná-Etendeka Province (133 Ma)
The Paraná-Etendeka Province was erupted in southern Gondwana (SE South America and Namibia) early in the Cretaceous (Renne et al. 1992), during the later half of the Valanginian Stage. Until recently, it was thought little of interest happened in the oceans at this time. However, recent studies have revealed an extremely watered-down version of the effects seen during the Early Jurassic and end-Permian crises. Thus, a thin, widespread, late Valanginian black shale has been found in ocean cores, indicating an episode of oxygen-poor deposition that has been called the Weissert Event (Erba et al. 2004). The calcareous nannoplankton fossil record shows a calcification crisis at the same time, which may reflect an increase in oceanic nutrient levels (calcareous nannoplankton are thought to prefer low-nutrient conditions) or acidification of ocean surface waters (Erba 2004) or both. This crisis did not cause extinctions and in fact proved something of a spur to evolution because planktonic foraminifera begin a Cretaceous-long radiation of new species immediately after the Weissert Event.
Unlike other intervals marked by LIP eruptions, it is unclear if any substantial global temperature changes occurred at this time. A case can be made for cooling in the last stages of the anoxic event, which probably reflects CO2-drawdown. This could be due to organic matter burial during the Weissert Event (Erba 2004), rather than be of volcanic origin. There is no negative carbon isotope anomaly associated with this event, and so methane release from gas hydrate is not likely.
Ontong Java Plateau (120 Ma)
A substantial part of the vast Ontong Java Plateau of the SW Pacific was erupted around the Barremian-Aptian boundary of the Early Cretaceous (Courtillot and Renne 2003). Eruption of this great volume of oceanic volcanic rocks slightly predated the oceanic anoxic event, known as the Selli Event, at 119 Ma (Larson and Erba 1999). However, just prior to this event there was a “nannoconid crisis” during which these very small, calcareous plankton suddenly became rather rare. This could reflect a calcification crisis, due to volcanogenic CO2 input, or a fertilisation crisis that did not favour the low nutrient-adapted nannoconids (Erba 2004). Erba further suggests that direct warming of the ocean water by the lava pile may have contributed to the break-down of ocean stratification and expansion of the oxygen-minimum zone. Both the warming and the fertilisation may have contributed to the anoxic event, but this effect was curiously delayed relative to the eruptions. The re-establishment of normal oceanic conditions after the Selli Event saw the reappearance of the missing nannoconids. Thus, the crisis was only temporary and not an extinction event.
Caribbean-Colombian Plateau (90 Ma)
The 133 and 119 Ma oceanic anoxic events appear to have precipitated only minor extinction crises when they are compared to the enormous losses of earlier mass extinctions. The next volcanism-anoxia event in the Cretaceous occurred at the Cenomanian-Turonian boundary, and this time it did coincide with rather more extinctions, notably of several planktonic foraminifera species (Wan et al. 2003). This interval also marks the culmination of Cretaceous greenhouse warming and sea level rise. It thus has many of the hallmarks of other volcanogenic crises, although there is only weak evidence for a calcification crisis and no evidence for methane release. This 90 Ma event coincided with the eruption of a LIP in the Caribbean-Colombian region and probably part of the Kerguelen LIP in the Indian Ocean, and also with some flood basalts in Madagascar (Kerr 1998). There were clearly a lot of volcanic culprits to choose from at this time.
Proposed kill mechanisms for the Cenomanian-Turonian extinctions include poisoning by trace metals derived from oceanic volcanism (Erba 2004), but this proposition is rather difficult to test. The anoxic event itself provides the most obvious cause of the marine extinctions, and the contribution of volcanism to global warming and fertilisation of the oceans provides a justification for linking volcanism and anoxia (Sinton and Duncan 1997). Furthermore, the oceanic volcanism may also have caused additional warming effects in addition to the direct input of volcanogenic CO2 to the atmosphere. Warming of the oceans by the lavas and oceanic acidification (by volcanic SO2 release) would both have released CO2 to the atmosphere, thus exacerbating a warming trend (Kerr this issue).
The Deccan Traps (65 Ma)
Evaluating the global environmental influence of the Deccan Trap eruptions in India is problematic due to the difficulty of disentangling the effects of the well-known, coeval Chicxulub impact event. However, thanks to prolonged and intensive study, the detailed chronology of volcanism and climate change in the Maastrichtian Stage, during the lead up to the end-Cretaceous mass extinction, are now established with some clarity. The mid-Maastrichtian was rather a cool interval, but a rapid phase of warming began around 400 kyr before the K-T boundary (Abramovich and Keller 2003). This was reversed by a rapid cooling trend around 100 kyr before the boundary, when the 4-5°C temperature gain was lost. The cooling coincides with a sharp sea level fall, and a lowstand was reached shortly before the K-T boundary. Thereafter, sea level began rising again across the boundary (Hallam and Wignall 1999).
These substantial oscillations in climate and sea level did not cause much in the way of extinctions. Keller (2003) has shown that the latest Maastrichtian warming pulse was associated with a destabilisation of planktonic foraminiferal populations and short-lived blooms of stress-tolerant species. According to Keller these may reflect the expansion and intensification of the mid-water oxygen-minimum zone. However, interesting though they are, these changes cannot compare with the near-total and abrupt mass extinction of planktonic foraminifera (and various other groups) at the end of the Cretaceous.
The possibility that the Deccan Trap eruptions were implicated in some or all of these changes has of course been known for some time. However, only recently has it been appreciated that the main eruptive phase coincided with the late Maastrichtian warm pulse (Ravizza and Peucker-Ehrenbrink 2003). The release of volcanic CO2 is the most likely driver of environmental change, with a calcification crisis in the oceans and global warming of the order of 4°C the most direct consequences. Thus, like the other LIP eruptions of the Cretaceous, the Deccan Trap eruptions appear to have caused significant climatic effects, but only modest biotic effects, perhaps because the oceans did not become anoxic. It has been argued that the biosphere was already rather stressed at the moment of meteorite impact, but without that impact one suspects the end-Maastrichtian event would only have ranked alongside minor Cretaceous crises such as the Selli Event (White and Saunders 2005).
North Atlantic Igneous (Brito-Arctic) Province (55 Ma)
The climatic events at 55 Ma, around the Palaeocene-Eocene (P-E) boundary, have received ample study and are reasonably well understood. Thus, a sharp negative δ13C excursion is generally taken as the signature of gas hydrate release, which in turn is held responsible for the brief (120 kyr) warming pulse at this boundary (Kennett and Stott 1991). Contemporary changes in the oceans include a calcification crisis and the development of oxygen-poor deep waters, which caused the extinction of many of the species living there. However, this was not a time of mass extinction by any stretch of the imagination. In fact extinction rates at this time were some of the lowest ever recorded.
These climatic and oceanic changes are very similar to the changes observed during Cretaceous LIP eruptions, and in this case they may relate to the eruption of the North Atlantic Igneous Province. However, this province seems to have been formed in two discrete pulses, with the younger pulse coinciding with the Palaeocene-Eocene thermal maximum at 55 Ma, and the older eruptive phase coinciding with a rather cool interval (Courtillot and Renne 2003). In a recent study of three marine P-E boundary sections, Schmitz et al. (2004) noted that the thermal maximum coincides with the onset of an unusual phase of intense, explosive basaltic volcanism in the North Atlantic region. Curiously the release of dust and aerosols should have produced cooling rather than the observed warming.
Perhaps the most intriguing question arising from the link between LIPs and environmental changes concerns the remarkably different magnitudes of the supposed volcanogenic effects. Thus, the Early Jurassic climatic and environmental changes are closely comparable to those proposed for the end-Permian crisis. Very similar changes also occurred during the Palaeocene-Eocene thermal maximum, but a mass extinction event has not been recorded. However, this last event was of much briefer duration and may not have lasted long enough to wreak the devastation of the earlier events.
In summary, large igneous province eruptions can cause changes that range from interesting but benign (Palaeocene-Eocene boundary), to severely damaging (Early Jurassic), to utterly catastrophic (end-Permian). A partial solution to this problem of variable influence may be found in modelling work. For example, Dessert et al. (2001) have suggested that factors such as pre-eruption atmospheric CO2 levels and the rate of eruption are key variables in any climatic changes. The closest correspondence between eruptions and extinctions coincides with the Pangean world, when most of the continents were part of a single supercontinent. It may be that such a configuration was less able to cope with sudden influxes of CO2 into the atmosphere because chemical weathering (the main mechanism of CO2 drawdown over geological timescales) would have been more limited in the arid interior of such a vast continent.
This paper has benefited from the comments of Gerta Keller and Andrew Kerr.