In order to classify volcanic eruptions and their potential effects on the atmosphere, Newhall and Self (1982) proposed a scale of explosive magnitude, the Volcanic Explosivity Index (VEI), based mainly on the volume of the erupted products (and the height of the volcanic eruption column). VEI’s range varies from VEI = 0 (for strictly non-explosive eruptions) to VEI = 8 (for explosive eruptions producing ~1012 m3 bulk volume of tephra). Eruption rates for VEI = 8 eruptions may be greater than 106 m3s−1 (Ninkovich et al., 1978a, 1978b).
Eruptions also differ in the amounts of sulphur-rich gases released to form stratospheric aerosols. Therefore, the sulphur content of the magma, the efficiency of degassing, and the heights reached by the eruption column are important factors in the climatic effects of eruptions (Palais and Sigurdsson, 1989; Rampino and Self, 1984). Historic eruptions of VEI ranging from 3 to 6 (volume of ejecta from <1 km3 to a few tens of km3) have produced stratospheric aerosol clouds up to a few tens of Mt. These eruptions, including Tambora 1815 and Krakatau 1883, have caused cooling of the Earth’s global climate of a few tenths of a degree Centigrade (Rampino and Self, 1984). The most recent example is the Pinatubo (Philippines) eruption of 1991 (Graf et al., 1993; Hansen et al., 1996).
Volcanic super-eruptions are defined as eruptions that are tens to hundreds of times larger than historic eruptions, attaining a VEI of 8 (Mason et al., 2004; Rampino, 2002; Rampino et al., 1988; Sparks et al., 2005). Super-eruptions are usually caldera-forming events and more than 20 super-eruption sites for the last 2 million years have been identified in North America, South America, Italy, Indonesia, the Philippines, Japan, Kamchatka, and New Zealand. No doubt additional super-eruption sites for the last few million years exist (Sparks et al., 2005).
The Late Pleistocene eruption of Toba in Sumatra, Indonesia was one of the greatest known volcanic events in the geologic record (Ninkovich et al., 1978a, 1978b; Rampino and Self, 1993a; Rose and Chesner, 1990). The relatively recent age and the exceptional size of the Toba eruption make it an important test case of the possible effects of explosive volcanism on the global atmosphere and climate (Oppenheimer, 2002; Rampino and Self, 1992, 1993a; Rampino et al., 1988; Sparks et al., 2005). For the Toba event, we have data on intercaldera fill, outflow sheets produced by pyroclastic flows and tephra fallout. Recent information on the environmental effects of super-eruptions supports the exceptional climatic impact of the Toba eruption, with significant effects on the environment and human population.
The Toba eruption has been dated by various methods K/Ar method at 73,500 ± 3500 yr BP (Chesner et al., 1991). The Toba ash layer occurs in deep-sea cores from the Indian Ocean and South China Sea (Huang et al., 2001; Shultz et al., 2002; Song et al., 2000). The widespread ash layer has a dense rock equivalent volume (DRE) of approximately 800 km3 (Chesner et al., 1991). The pyroclastic flow deposits on Sumatra have a volume of approximately 2000 km3DRE (Chesner et al., 1991; Rose and Chesner, 1990), for a total eruption volume of approximately 2800 km3 (DRE). Woods and Wohletz (1991) estimated Toba eruption cloud heights of 32 ± 5 km, and the duration of continuous fallout of Toba ash over the Indian Ocean has been estimated at two weeks or less (Ledbetter and Sparks, 1979).
Release of sulphur volatiles is especially important for the climatic impact of an eruption, as these form sulphuric acid aerosols in the stratosphere (Rampino and Self, 1984). Although the intrinsic sulphur content of rhyolite magmas is generally low, the great volume erupted is sufficient to give an enormous volatile release. Based on studies of the sulphur content of the Toba deposits, Rose and Chesner (1990) estimated that approximately 3 × 1015 g of H2S/SO2 (equivalentto ~ 1 × 1016 g of H2SO4 aerosols) could have been released from the erupted magma. The amounts of fine ash and sulphuric acid aerosols that could have been generated by Toba was estimated independently using data from smaller historical rhyolitic eruptions (Rampino and Self, 1992). By this simple extrapolation, the Toba super-eruption could have produced up to 2 × 1016 g of fine (<2 μ) dust and approximately 1.5 × 1015 g of sulphuric acid aerosols.
Physical and chemical processes in dense aerosol clouds may act in a ‘self-limiting’ manner, significantly reducing the amount of long-lived H2SO4 aerosols (Rampino and Self, 1982; Pinto et al., 1989). Using one-dimensional aerosol microphysical and photochemical models, Pinto and others (1989) showed that for an aerosol cloud of approximately 1014 g of SO2, condensation and coagulation are important in producing larger-sized particles, which have a smaller optical effect per unit mass, and settle out of the stratosphere faster than smaller particles. However, the maximum sulphur volatile emission that they modelled was 2 × 1014 g of SO2, and no data exist on the behaviour of H2SO4 aerosols in more than 10 times denser clouds.
Another possible limitation on aerosol loading is the amount of water in the stratosphere available to convert SO2 to H2SO4. Stothers et al. (1986) calculated that approximately 4 × 1015 g of water might be available in the ambient stratosphere, and injection into the stratosphere of up to 5.4 × 1017 g of H2O from Toba is possible (Rose and Chesner, 1990), more than enough water to convert the sulphur gases emitted by Toba into H2SO4 aerosols.
The exceptional magnitude of the Toba eruption makes it a natural target in the studies of large volcanic events preserved in polar ice cores. Work on the GISP2 ice core from Summit, Greenland, revealed an approximately 6-year long period of enhanced volcanic sulphate dated at 71, 100 ± 5000 years ago identified with the Toba eruption (Zielinski et al., 1996a, 1996b). The magnitude of this sulphate signal is the largest in the entire 110,000 years of the GISP2 record.
Zielinski and others (1996a) estimated that the total atmospheric loading of H2SO4 for the approximately 6-year period of the ice-core peak ranged from approximately 0.7 to 4.4 × 1015 g, in general agreement with the above estimates derived from volcano-logical techniques and scaling from smaller eruptions (Rampino and Self, 1992, 1993a; Rose and Chesner, 1990). Estimates of aerosol loadings range from approximately 150 to 1000 Mt per year, over the approximately 6-year period of the ice-core peak.
The SO2-4 signal identified with Toba coincides with the beginning of an approximately 1000-year cooling event seen in the ice-core record between brief warm periods (interstadials), but is separated from the most recent major approximately 9000-year glacial period by the approximately 2000-year-long warmer period. A similar cool pulse between interstadials is seen in the pollen record of the Grande Pile in northeastern France, dated as approximately 70,000 years BP (Woillard and Mook, 1982).
Thus, the ice-core evidence suggests that the Toba signal occurred during the transition from a warm interglacial climate and was preceded and followed by abrupt climate oscillations that preceded the start of the most recent major early glaciation (Zielinski et al., 1996a, 1996b).
Since Toba is a low-latitude volcano, dust and volatiles would have been injected efficiently into both Northern and Southern Hemispheres (Rampino et al., 1988), although the season of the eruption is unknown. These estimated aerosol optical effects are roughly equivalent in visible opacity to smoke-clouds (Turco et al., 1990), which is within the range used in nuclear-winter scenarios of massive emissions of soot emanating from burning urban and industrial areas in the aftermath of nuclear war.
Although the climate conditions and duration of a nuclear winter have been much debated, simulations by Turco and others (1990) predicted that land temperatures in the 30°-70°N latitude zone could range from approximately 5°C to approximately 15°C colder than normal, with freezing events in mid-latitudes during the first few months. At lower latitudes, model simulations suggest cooling of 10°C or more, with drastic decreases in precipitation in the first few months. Ocean-surface cooling of approximately 2–6°C might extend for several years, and persistence of significant soot for 1–3 years might lead to longer term (decadal) climatic cooling, primarily through climate feedbacks including increased snow cover and sea ice, changes in land surface albedo, and perturbed sea-surface temperatures (Rampino and Ambrose, 2000).
The injection of massive amounts of volcanic dust into the stratosphere by a super-eruption such as Toba might be expected to lead to similar immediate surface cooling, creating a ‘volcanic winter’ (Rampino and Self, 1992; Rampino et al., 1988). Volcanic dust probably has a relatively shorter residence time in the atmosphere (3-6 months) than soot (Turco et al., 1990) and spreads from a point source, but volcanic dust is injected much higher into the stratosphere, and hence Toba ash could have had a wide global coverage despite its short lifetime. Evidence of the wide dispersal of the dust and ash from Toba can be seen from lake deposits in India, where the reworked Toba ash forms a layer up to 3 m thick, and from the widespread ash layer in the Indian Ocean and South China Sea (Acharya and Basu, 1993; Huang et al., 2001; Shane et al., 1995).
Evidence for rapid and severe cooling from the direct effects of volcanic ash clouds comes from the aftermath of the 1815 Tambora eruption. Madras, India experienced a dramatic cooling during the last week of April 1815, a time when the relatively fresh ash and aerosol cloud from Tambora (10-11 April) would have been overhead. Morning temperatures dropped from 11°C on Monday to – 3°C on Friday (Stothers, 1984a). A similar, but much smaller effect, occurred as the dust cloud from the 1980 Mt St Helens eruption passed over downwind areas (Robock and Mass, 1982).
The stratospheric injection of sulphur volatiles (>1015 g), and the time required for the formation and spread of volcanic H2SO4 aerosols in the stratosphere should lead to an extended period of increased atmospheric opacity and surface cooling. The ice-core record, however, indicates stratospheric loadings of 1014 to 1015 g of H2SO4 aerosols for up to 6 years after the eruption (Zielinski et al., 1996a).
This agrees with model calculations by Pope and others (1994) that predict oxidation lifetimes (time required to convert a given mass of sulphur into H2SO4 aerosols) of between 4 and 17 years, and diffusion lifetimes (time required to remove unoxidized SO2 by diffusion to the troposphere) of between 4 and 7 years for total sulphur masses between 1015 and 1016 g. For atmospheric injection in this range, the diffusion lifetime is the effective lifetime of the cloud because the SO2 reservoir is depleted before oxidation is completed.
If the relationship between Northern Hemisphere cooling and aerosol loading from large eruptions is approximately linear, then scaling up from the 1815 AD. Tambora eruption would lead to an approximately 3.5°C hemispheric cooling after Toba (Rampino and Self, 1993a). Similarly, empirical relationships between SO2 released and climate response (Palais and Sigurdsson, 1989) suggested a hemispheric surface-temperature decrease of about 4 ± 1°C. The eruption clouds of individual historic eruptions have been too short-lived to drive lower tropospheric temperatures to their steady-state values (Pollack et al., 1993), but the apparently long-lasting Toba aerosols may mean that the temperature changes in the troposphere attained a larger fraction of their steady-state values. Huang et al. (2001) were able to correlate the Toba ash in the South China Sea with a 1°C cooling of surface waters that lasted about 1000 years.
Considering a somewhat smaller super-eruption, the Campanian eruption of approximately 37,000 cal yr BP in Italy (150 km3 of magma discharged) was coincident with Late Pleistocene bio-cultural changes that occurred within and outside the Mediterranean region. These included the Middle to Upper Paleolithic cultural transition and the replacement of Neanderthals by ‘modern’ Homo sapiens (Fedele et al., 2002).
The climatic and environmental impacts of the Toba super-eruption are potentially so much greater than that of recent historical eruptions (e.g., Hansen et al., 1992; Stothers, 1996) that instrumental records, anecdotal information, and climate-model studies of the effects of these eruptions may not be relevant in scaling up to the unique Toba event (Rampino and Self, 1993a; Rampino et al., 1988). Various studies on the effects of extremes of atmospheric opacity and climate cooling on the environment and life have been carried out, however, in connection with studies of nuclear winter and the effects of asteroid impacts on the earth (e.g., Green et al., 1985; Harwell, 1984; Tinus and Roddy, 1990), and some of these may be relevant to the Toba situation.
Two major effects on plant life from high atmospheric opacity are reduction of light levels and cold temperatures. Reduction in light levels expected from the Toba eruption would range from dim-sun conditions (~75% sunlight transmitted) like those seen after the 1815 Tambora eruption, to that of an overcast day (~10% sunlight transmitted). Experiments with young grass plants have shown how net photosynthesis varies with light intensity. For a decrease to 10% of the noon value for a sunny summer day, photosynthesis was reduced by about 85% (van Kuelan et al., 1975), and photosynthesis also drops with decreasing temperatures (Redman, 1974).
Resistance of plants to unusual cold conditions varies somewhat. Conditions in the tropical zone are most relevant to possible impacts on early human populations in Africa. Tropical forests are very vulnerable to chilling, and Harwell and others (1985) argue that for freezing events in evergreen tropical forests, essentially all aboveground plant tissues would be killed rapidly.
Average surface temperatures in the tropics today range from approximately 16–24°C. Nuclear winter scenarios predict prolonged temperature decreases of 3–7°C in Equatorial Africa, and short-term temperature decreases of up to 10°C. Many tropical plants are severely damaged by chilling to below 10–15°C for a few days (Greene et al., 1985; Leavitt, 1980). Most tropical forest plants have limited seed banks, and the seeds typically lack a dormant phase. Furthermore, regrowth tends to produce forests of limited diversity, capable of supporting much less biomass (Harwell et al., 1985).
Even for temperate forests, destruction could be very severe (Harwell, 1984; Harwell et al., 1985). In general, the ability of well-adapted trees to withstand low temperatures (cold hardiness) is much greater than that needed at any single time of the year, but forests can be severely damaged by unusual or sustained low temperatures during certain times of the year. A simulation of a 10°C decrease in temperatures during winter shows a minimal effect on the cold-hardy and dormant trees, whereas a similar 10°C drop in temperature during the growing season (when cold hardiness is decreased) leads to a 50% dieback, and severe damage to surviving trees, resulting in the loss of at least a year’s growth.
The situation for deciduous forest trees would be even worse than that for the evergreens, as their entire foliage would be new and therefore lost. For example, Larcher and Bauer (1981) determined that cold limits of photosynthesis of various temperate zone plants range from –1.3 to –3.9°C, approximately the same range as the tissue-freezing temperatures for these plants. Lacking adequate food reserves, most temperate forest trees would not be able to cold harden in a timely manner, and would die or suffer additional damage during early freezes in the Fall (Tinus and Roddy, 1990).
The effect of the Toba super-eruption on the oceans is more difficult to estimate. Regionally, the effect on ocean biota of the fallout of approximately 4 g/cm2 of Toba ash over an area of 5 × 106 km2 in the Indian Ocean must have been considerable. Deposition rates of N, organic C, and CaCO3 all rise sharply in the first few centimetres of the Toba ash layer, indicating that the ash fallout swept the water column of most of its particulate organic carbon and calcium carbonate (Gilmour et al., 1990).
Another possible effect of a dense aerosol cloud is decreased ocean productivity. For example, satellite observations after the 1982 El Chichón eruption showed high aerosol concentrations over the Arabian Sea, and these values were associated with low surface productivity (as indicated by phytoplankton concentrations) from May through October of that year (Strong, 1993). Brock and McClain (1992) suggested that the low productivity was related to weaker-than-normal monsoon winds, and independent evidence suggests that the southwest monsoon in the area arrived later and withdrew earlier than usual, and that the wind-driven Somali current was anomalously weak. Conversely, Genin and others (1995) reported enhanced vertical mixing of cooled surface waters in weakly stratified areas of the Red Sea following the Pinatubo eruption, which resulted in algal and phytoplankton blooms that precipitated widespread coral death.
Studies following the 1991 Pinatubo eruption provide evidence that aerosol-induced cooling of the southwestern Pacific could lead to significant weakening of Hadley Cell circulation and rainfall, and might precipitate long-term El Niño-like anomalies with extensive drought in many tropical areas (Gagan and Chivas, 1995). Some climate-model simulations predict significant drought in tropical areas from weakening of the trade winds/Hadley circulation and from reduction in the strength of the summer monsoon (e.g., Pittock et al., 1986, 1989; Turco et al., 1990). For example, Pittock and others (1989) presented GCM results that showed a 50% reduction in convective rainfall in the tropics and monsoonal regions.
Recent debate about the origin of modern humans has focused on two competing hypotheses: (1) the ‘multiregional’ hypothesis, in which the major subdivisions of our species evolved slowly and in situ, with gene flow accounting for the similarities now observed among groups, and (2) the ‘replacement’ hypothesis, in which earlier populations were replaced 30,000 to 100,000 years ago by modern humans that originated in Africa (Hewitt, 2000; Rogers and Joude, 1995).
Genetic studies have been used in attempts to test these two hypotheses. Studies of nuclear and mitochondrial DNA from present human populations led to the conclusion that the modern populations originated in Africa and spread to the rest of the Old World approximately 50,000 ± 20,000 years ago (Harpending et al., 1993; Jones and Rouhani, 1986; Wainscoat et al., 1986). This population explosion apparently followed a severe population bottleneck, estimated by Harpending and others (1993) to have reduced the human population to approximately 500 breeding females, or a total population as small as 4000 for approximately 20,000 years. At the same time, Neanderthals who were probably better adapted for cold climate moved into the Levant region when modern humans vacated it (Hewitt, 2000).
Harpending and others (1993) proposed that the evidence may fit an intermediate ‘Weak Garden of Eden’ hypothesis that a small ancestral human population separated into partially isolated groups about 100,000 years ago, and about 30,000 years later these populations underwent either simultaneous bottlenecks or simultaneous expansions in size. Sherry and others (1994) estimated mean population expansions times ranging from approximately 65,000 to 30,000 years ago, with the African expansion possibly being the earliest.
Ambrose (1998, 2003; see Gibbons, 1993) pointed out that the timing of the Toba super-eruption roughly matched the inferred timing of the bottleneck and release, and surmised that the environmental after-effects of the Toba eruption might have been so severe as to lead to a precipitous decline in the population of human ancestors (but see Gathorne-Hardy and Harcourt-Smith, 2003 for opposing views). Rampino and Self (1993a) and Rampino and Ambrose (2000) concurred that the climatic effects of Toba could have constituted a true ‘volcanic winter’, and could have caused severe environmental damage. It may be significant that analysis of mtDNA of Eastern Chimpanzee (Pan troglodytes schweinfurthii) shows a similar pattern to human DNA, suggesting a severe reduction in population at about the same time as in the human population (see Rogers and Jorde, 1995).
Decker (1990) proposed that if all magnitude 8 eruptions in the recent past left caldera structures that have been recognized, then the frequency of VEI 8 eruptions would be approximately 2 × 10 – 5 eruptions per year, or roughly one VEI 8 eruption every 50,000 years.
The timing and magnitude of volcanic eruptions, however, are difficult to predict. Prediction strategies have included (1) recognition of patterns of eruptions at specific volcanoes (e.g., Godano and Civetta, 1996; Klein, 1982), (2) precursor activity of various kinds (e.g., Chouet, 1996; Nazzaro, 1998), (3) regional and global distribution of eruptions in space and time (Carr, 1977; Mason et al., 2004; Pyle, 1995), and (4) theoretical
predictions based on behaviour of materials (Voight, 1988; Voight and Cornelius, 1991). Although significant progress has been made in short-term prediction of eruptions, no method has proven successful in consistently predicting the timing, and more importantly, the magnitude of the resulting eruption or its magmatic sulphur content and release characteristics. State-of-the-art technologies involving continuous satellite monitoring of gas emissions, thermal anomalies and ground deformation (e.g., Alexander, 1991; Walter, 1990) promise improved forecasting and specific prediction of volcanic events, but these technologies are thus far largely unproven.
For example, although we have 2000 years of observations for the Italian volcano Vesuvius (Nazzaro, 1998), and a long history of monitoring and scientific study, prediction of the timing and magnitude of the next Vesuvian eruption remains a problem (Dobran et al., 1994; Lirer et al., 1997). For large caldera-forming super-eruptions, which have not taken place in historic times, we have little in the way of meaningful observations on which to base prediction or even long-range forecasts.
The regional and global effects of the ash fallout and aerosol clouds on climate, agriculture, health, and transportation would present a severe challenge to modern civilization. The major effect on civilization would be through collapse of agriculture as a result of the loss of one or more growing seasons (Toon et al., 1997). This would be followed by famine, the spread of infectious diseases, breakdown of infrastructure, social and political unrest, and conflict. Volcanic winter predictions are for global cooling of 3–5°C for several years, and regional cooling up to 15°C (Rampino and Self, 1992; Rampino and Ambrose, 2000). This could devastate the major food-growing areas of the world. For example, the Asian rice crop could be destroyed by a single night of below-freezing temperatures during the growing season. In the temperate grain-growing areas, similar drastic effects could occur. In Canada, a 2–3°C average local temperature drop would destroy wheat production, and 3–4°C would halt all Canadian grain production. Crops in the American Midwest and the Ukraine could be severely injured by a 3–4°C temperature decrease (Harwell and Hutchinson, 1985; Pittock et al., 1986). Severe climate would also interfere with global transportation of foodstuffs and other goods. Thus, a super-eruption could compromise global agriculture, leading to famine and possible disease pandemics (Stothers, 2000).
Furthermore, large volcanic eruptions might lead to longer term climatic change through positive feedback effects on climate such as cooling the surface oceans, formation of sea-ice, or increased land ice (Rampino and Self, 1992, 1993a, 1993b), prolonging recovery from the ‘volcanic winter’. The result could be widespread starvation, famine, disease, social unrest, financial collapse, and severe damage to the underpinnings of civilization (Sagan and Turco, 1990; Sparks et al., 2005).
The location of a super-eruption can also be an important factor in its regional and global effects. Eruptions from the Yellowstone Caldera over the last 2 million years have included three super-eruptions. Each of these produced thick ash deposits over the western and central United States (compacted ash thicknesses of 0.2 m occur ~1500 km from the source; Wood and Kienle, 1990).
One mitigation strategy could involve the stockpiling of global food reserves. In considering the vagaries of normal climatic change, when grain stocks dip below about 15% of utilization, local scarcities, worldwide price jumps and sporadic famine were more likely to occur. Thus a minimum world level of accessible grain stocks near 15% of global utilization should be maintained as a hedge against year-to-year production fluctuations due to climatic and socioeconomic disruptions. This does not take into account social and economic factors that could severely limit rapid and complete distribution of food reserves.
At present, a global stockpile equivalent to a 2-month global supply of grain exists, which is about 15% of annual consumption. For a super-volcanic catastrophe, however, several years of growing season might be curtailed, and hence a much larger stockpile of grain and other foodstuffs would have to be maintained, along with the means for rapid global distribution.
The chances for communicative intelligence in the Galaxy is commonly represented by a combination of the relevant factors called the Drake Equation, which can be written as
where N is the number of intelligent communicative civilizations in the Galaxy; R* is the rate of star formation averaged over the lifetime of the Galaxy; fp is the fraction of stars with planetary systems; ne is the mean number of planets within such systems that are suitable for life; fl is the fraction of such planets on which life actually occurs; fi is the fraction of planets on which intelligence of arises; fc is the fraction of planets on which intelligent life develops a communicative phase; and L is the mean lifetime of such technological civilizations (Sagan, 1973).
Although the Drake Equation is useful in organizing the factors that are thought to be important for the occurrence of extraterrestrial intelligence, the actual assessment of the values of the terms in the equation is difficult. The only well-known number is R*, which is commonly taken as 10 yr – 1. Estimates for N have varied widely from approximately 0 to >108 civilizations (Sagan, 1973).
It has been pointed out recently that fc and L are limited in part by the occurrence of asteroid and comet impacts that could prove catastrophic to technological civilizations (Sagan and Ostro, 1994; Chyba, 1997). Present human civilization, dependent largely on annual crop yields, is vulnerable to an ‘impact winter’ that would result from dust lofted into the stratosphere by the impact of objects ≥1 km in diameter (Chapman and Morrison, 1994; Toon et al., 1997). Such an impact would release approximately 105-106 Mt (TNT equivalent) of energy, produce a crater approximately 20–40 km in diameter, and is calculated to generate a global cloud consisting of approximately 1000 Mt of submicron dust (Toon et al., 1997). Covey et al. (1990) performed 3-D climate-model simulations for a global dust cloud containing submicron particles with a mass corresponding that that produced by an impact of 6 × 105 Mt (TNT). In this model, global temperatures dropped by approximately 8 C during the first few weeks. Chapman and Morrison (1994) estimated that an impact of this size would kill more than 1.5 billion people through direct and non-direct effects.
Impacts of this magnitude are expected to occur on average about every 100,000 years (Chapman and Morrison, 1994). Thus, a civilization must develop science and technology sufficient to detect and deflect such threatening asteroids and comets on a time scale shorter than the typical times between catastrophic impacts. Recent awareness of the impact threat to civilization has led to investigations of the possibilities of detection, and deflection or destruction of asteroids and comets that threaten the Earth (e.g., Gehrels, 1994; Remo, 1997). Planetary protection technology has been described as essential for the long-term survival of human civilization on the Earth.
The drastic climatic and ecological effects predicted for explosive super-eruptions leads to the question of the consequences for civilization here on Earth, and on other earth-like planets that might harbour intelligent life (Rampino, 2002; Sparks et al., 2005). Chapman and Morrison (1994) suggested that the global climatic effects of super-eruptions such as Toba might be equivalent to the effects of an approximately 1 km diameter asteroid. Fine volcanic dust and sulphuric acid aerosols have optical properties similar to the submicron dust produced by impacts (Toon et al., 1997), and the effects’ on atmospheric opacity should be similar. Volcanic aerosols, however, have a longer residence time of several years (Bekki et al., 1996) compared to a few months for fine dust, so a huge eruption might be expected to have a longer lasting effect on global climate than an impact producing a comparable amount of atmospheric loading.
Estimates of the frequency of large volcanic eruptions that could cause ‘volcanic winter’ conditions suggest that they should occur about once every 50,000 years. This is approximately a factor of two more frequent than asteroid or comet collisions that might cause climate cooling of similar severity (Rampino, 2002). Moreover, predicting or preventing a volcanic climatic disaster might be more difficult than tracking and diverting incoming asteroids and comets. These considerations suggest that volcanic super-eruptions pose a real threat to civilization, and efforts to predict and mitigate volcanic climatic disasters should be contemplated seriously (Rampino, 2002; Sparks et al., 2005).
Acknowledgement
I thank S. Ambrose, S. Self, R. Stothers, and G. Zielinski for the information provided.
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