Large igneous provinces and mass extinctions

Ž .Earth-Science Reviews 53 2001 1–33www.elsevier.comrlocaterearscirev

Large igneous provinces and mass extinctions

P.B. WignallDepartment of Earth Sciences, UniÕersity of Leeds, Leeds LS2 9JT, UK

Abstract

Comparing the timing of mass extinctions with the formation age of large igneous provinces reveals a closecorrespondence in five cases, but previous claims that all such provinces coincide with extinction events are undulyoptimistic. The best correlation occurs for four consecutive mid-Phanerozoic examples, namely the end-GuadalupianextinctionrEmeishan flood basalts, the end-Permian extinctionrSiberian Traps, the end-Triassic extinctionrcentral Atlanticvolcanism and the early Toarcian extinctionrKaroo Traps. Curiously, the onset of eruptions slightly post-dates the mainphase of extinctions in these examples. Of the seven post-Karoo provinces, only the Deccan Traps coincide with a massextinction, but in this case, the nature of the biotic crisis is best reconciled with the effects of a major bolide impact.Intraoceanic volcanism may also be implicated in a relatively minor end-Cenomanian extinction crisis, although once againthe main phase of volcanism occurs after the crisis. The link between large igneous province formation and extinctionsremains enigmatic; volume of extrusives and extinction intensity are unrelated and neither is there any apparent relationship

Ž .with the rapidity of province formation. Violence of eruptions proportions of pyroclastics also appears unimportant. Six outof 11 provinces coincide with episodes of global warming and marine anoxiardysoxia, a relationship that suggests thatvolcanic CO emissions may have an important effect on global climate. Conversely, there is little, if any, geological2

evidence for cooling associated with continental flood basalt eruptions suggesting little long-term impact of SO emissions.2

Large carbon isotope excursions are associated with some extinction events and intervals of flood basalt eruption but theseare too great to be accounted for by the release of volcanic CO alone. Thus, voluminous volcanism may in some2

Ž .circumstances trigger calamitous global environmental changes runaway greenhouses , perhaps by causing the dissociationof gas hydrates. The variable efficiency of global carbon sinks during volcanic episodes may be an important control onenvironmental effects and may explain why the eruption of some vast igneous provinces, such as the Parana–Etendeka´Traps, have little perceptible climatic impact. q 2001 Elsevier Science B.V. All rights reserved.

Keywords: mass extinctions; flood basalts; marine anoxia; global warming; runaway greenhouse

1. Introduction

The notion that volcanicity is capable of causingthe global devastation required to cause mass extinc-tions is a relatively recent development in the scien-

Ž .tific literature. Vogt 1972 was the first to note thatthe eruption of the Deccan Traps flood basaltprovince roughly coincided with the end-Cretaceousmass extinction, which led him to speculate that

trace metals released by plume eruptions may havecaused extinction by poisoning. Other early worksuggested that the cooling effect of volcanic ash andsulphate aerosols injected into the stratosphere dur-ing large Plinian eruptions may also have been suffi-

Žciently intense to cause mass extinctions Budyko.and Pivivariva, 1967; Axelrod, 1981 , while Pollock

Ž .et al. 1976 suggested that a series of closely spacederuptions may be capable of triggering an ice age.

0012-8252r01r$ - see front matter q 2001 Elsevier Science B.V. All rights reserved.Ž .PII: S0012-8252 00 00037-4

( )P.B. WignallrEarth-Science ReÕiews 53 2001 1–332

However, the debate on volcanically caused extinc-tion only truly began with the seminal paper of

Ž .Alvarez et al. 1980 , and their proposition that theend-Cretaceous mass extinction was caused by bolideimpact. Deccan Traps volcanicity rapidly becameestablished as the other principal contender for the

Žcause of this event McLean, 1981; Officer andDrake, 1983; Officer et al., 1987; Courtillot and

.Besse, 1987 . Subsequent debate has focussed onthis and the other continental flood basalt provincesŽ . Ž .CFBPs and large igneous provinces LIPs in gen-eral. The likelihood of a cause-and-effect link be-tween mass extinctions and CFBPsrLIPs has beenconsiderably strengthened by the discovery of theclose temporal relationship between these phenom-

Žena Rampino and Stothers, 1988; Stothers andRampino, 1990; Stothers, 1993; Courtillot, 1994;

.Hallam and Wignall, 1997; Olsen, 1999 .The precise nature of this correspondence is an

intriguing one which forms the central part of thisreview article. The potential links between LIPs andmass extinctions are examined on a case-by-casebasis with emphasis given to the crucial aspect ofrelative timing. Evidence for volcanically triggeredenvironmental changes during extinction crises arealso reviewed in the light of the potential climaticeffects of volcanism discussed below. It is importantto note that mass extinctions are generally defined as

Ž .geologically brief intervals i.e. ;1 Ma of elevatedextinction rates that affected diverse taxa from a

Ž .broad range of habitats i.e. terrestrial and marineŽ .throughout the world Hallam and Wignall, 1997 . In

practise extinction events are identified as peaks inthe extinction rate-versus-time graphs of SepkoskiŽ .1982, 1996 . Only the five biggest peaks on thesegraphs are widely regarded as true mass extinctions,although several lesser extinction peaks coincide withLIP formation.

2. Climatic effects of volcanicity

2.1. Short-term effects

Volcanoes release a range of gases into the strato-sphere of which SO and CO are volumetrically the2 2

Ž .most important Sigurdsson, 1990a,b . Their effectsare diametrically opposed and operate over different

timescales. Sulphur dioxide is a greenhouse gas andits initial effect is to cause warming, however, itreacts rapidly with water in the atmosphere to pro-duce sulphate aerosols that backscatter and absorb

Žthe sun’s radiation Devine et al., 1984; Sigurdsson,.1990a,b; Fig. 1 . Such effects are localised to the

vicinity of the eruption unless the gases are injectedinto the lower stratosphere whereupon they arerapidly dispersed around the hemisphere. Globalcooling from Plinian-style eruptions has been well

Ž .recorded in historical times e.g. Genin et al., 1995 ,although the effect is usually only of 1 to 2 yearsduration due to the rapid rain-out of the sulphate

Ž .aerosol Pinto et al., 1989; Officer et al., 1987 .Volcanic ash remains in the atmosphere for evenshorter durations and so is unlikely to contribute

Ž .significantly to cooling Devine et al., 1984 . Frominterpretation of historical records, Devine et al.Ž .1984 discovered a roughly geometric relationshipbetween mass of S released and northern hemisphere

Ž .temperature decline Fig. 2 . This suggests that largervolcanic eruptions than those witnessed in historical

Ž .times may have caused substantial but short-livedcooling. However, the connection between eruptionsand cooling is far from straightforward: several ofthe best documented historical connections may be

Ž .no more than fortuitous Mass and Portman, 1989 .For example, the cold summer of 1816 in NorthAmerica has been attributed to the Toba eruption theprevious year. However, 1816 was just one of manycold summers of the early 19th century and it was by

Ž .no means the most severe Sadler and Grattan, 1999 .The largest explosive eruption of the recent geo-

logical past also occurred at Toba approximately 73Ž .ka; Rampino and Self 1992 estimated that the

resultant ash layer contains 800 km3 of magma.Ž .Extrapolating from the graph of Devine et al. 1984 ,

they estimated that this eruption should have pro-duced 3.58C cooling which, they postulate, may havetriggered an ice age. However, as Rampino and SelfŽ .1992 noted, the transition from interglacial toglacial conditions occurred at 74 ka, slightly beforethe Toba eruption. New discoveries suggest that theToba ash band was even more widespread than

Ž .Rampino and Self 1992 estimated, and probablyrepresents greater than 3000 km3 of tephra. Zielinski

Ž .et al. 1996 calculate that such an enormous erup-tion is likely to have injected between 2200 and

( )P.B. WignallrEarth-Science ReÕiews 53 2001 1–33 3

Fig. 1. Effects of volcanic gases and the intervals over which they operate. With the exception of CO , most gases are rapidly removed from2

the atmosphere and are thus able to effect weather rather than long-term climate.

4400 Mt of sulphate aerosols into the stratosphereand their discovery of elevated sulphate concentra-tions in 6 years-worth of ice in the GISP2 core fromEast Greenland provides a valuable clue to the dura-tion of the potential climatic effects. However, the

Fig. 2. Hemispheric cooling effect of volcanic SO emissions2Žbased on inferred effects of historical eruptions from Devine et

.al., 1984 .

detailed d18 O climatic record for the interval imme-

diately preceding the sulphate-rich ice indicates 800years of cooling with a transition to a 1000-yearstadial beginning just a few years before the rise insulphate values. These records therefore lend nosupport to the notion that even the largest volcaniceruptions are capable of causing climatic change,

Ž .although Zielinski et al. 1996 suggest that thestadial interval may have been prolonged because ofthe Toba eruption.

Records of gigantic volcanic eruptions have alsoŽ .been documented by Huff et al. 1992 from the

Ordovician of North America. They studied K-be-Ž .ntonites from the mid-Caradoc Series ;454 Ma

and showed that the thickest example, which can betraced to Scandanavia, probably contains over 1000km3 of magma. Like the 73 ka Toba event, there isno discernible climatic change or extinction event

Ž .associated with this eruption Huff et al., 1992 .Reviewing other major pyroclastic eruptions, Erwin

Ž .and Vogel 1992 similarly concluded that they havenot caused extinctions.

The apparently benign influence of aerosol-in-duced volcanic winters may relate to self-limiting


effects and the high thermal heat capacity and thusthe thermal inertia of the oceans. Even temperaturedeclines of 3–58C occurring over a few years are

Žunlikely to change sea surface temperatures Loper.et al., 1988 , although the effects of such cooling

may be more severe for terrestrial environments.Ž .Modelling experiments of Pinto et al. 1989 have

also shown that as the rate of SO release into the2

atmosphere increases, progressively larger aerosolŽparticles are formed rather than more of the same

.size which settle out at a faster rate. However, theamount of water vapour in the atmosphere during theeruption interval may become limiting during largervolcanic events with the result that sulphate aerosolsform at a slower rate. This may explain the 5–6 yearduration of sulphate-rich record in the GISP2 icecore. Thus, the cooling effect of the largest volcaniceruptions may be more prolonged than for smallereruptions but not necessarily any more intense.

Of the other volcanic gases, only Cl has beenpostulated to occur in sufficient concentrations tocause environmental harm. Potential damage in-cludes the localised effects of acid rain and ozone

Ž .depletion Cockell, 1999 . However, the Acold trapeffectB, in which HCl and water vapour condense onash particles, provides a mechanism for the rapid

Žremoval of volcanic Cl from the atmosphere Pinto.et al., 1989 .

2.2. Long-term climatic effects

Volcanic activity also injects CO into the atmo-2

sphere and, unlike SO , this cannot be rapidly re-2Ž .moved Fig. 1 . However, the estimated annual CO2

input from volcanism is roughly 1011 kg, a figurethat is dwarfed by the current anthropogenic CO2

release into the atmosphere of 1013 kg yeary1

Ž .Leavitt, 1982; Sigurdsson 1990a , and both valuesare tiny compared to the 5=1019 g C in the atmo-sphere and oceans. Thus, single volcanic eruptionsare unlikely to cause any noticeable increase in thisimportant greenhouse gas. However, because of thelong residence time of CO in the surficial system,2

the cumulative effects of successive large eruptions,typical of those encountered in CFBPs, could beclimatically significant. Calculation of the amount ofCO released during eruptions can be obtained from2

Ž .Leavitt’s 1982 empirically derived formula. Thus,

the number of moles of CO released during a2

volcanic eruption is given by:

m CO ssf Õdrm.w. CO2 dg 2

where s is the weight fraction of CO in magma, f2 dg

is the fraction of gas evolved from the magma, Õ isthe magma volume, d is the solid density of themagma and m.w. CO is the molecular weight of2

Ž .CO 44 g . For basaltic lavas typical values of s, f2 dg

and d are 0.002, 0.6 and 2.9=1015 g kmy3, respec-Ž .tively Leavitt, 1982 .

2.3. Effects of flood basalt fissure eruptions

This brief review of historical and geologicalvolcanism provides little evidence for a link betweenindividual volcanic eruptions and climate change,and no evidence at all for a link with extinctions.However, the volumetrically most important volcanicactivity consists of the largely non-explosive erup-tion of flood basalts from fissures: such flows com-prise the bulk of LIPs and none have occurred in thelast few million years. Basaltic lavas are substan-tially more SO -rich than their acidic equivalents2Ž .Sigurdsson, 1982 , potentially indicating that theglobal cooling effect of sulphate aerosols may besevere during voluminous basalt eruptions. Basaltlavas are, however, relatively poor in halogen gaseshaving less than 50% the Cl content and 60% of the

Ž .F content of acidic lavas Sigurdsson, 1990b .The only substantial example of a basalt fissure

eruption in historical times occurred at Laki in Ice-Ž .land between 1783 and 1784 Sigurdsson, 1982 . A

total of 14.7 km3 of lava quietly flowed from a25-km-long fissure, over 60% of it in the first 40days of the eruption, and released an estimated 122

ŽMt of SO , 15 Mt of HF and 7 Mt of HCl Thordar-2.son et al., 1996 . The local effects were devastating

Žwith a reported dark bluish haze probably a sulphate.aerosol blanketing Iceland and causing crop failure

Ž .and famine Sigurdsson, 1982 . Contemporaryrecords indicate the summer of 1783 was exception-ally warm in northern Europe, perhaps a conse-

Žquence of the greenhouse effect of SO Grattan and2.Sadler, 1999 . The following winter of 1783–84 was

unusually cold in the northern hemisphere, as notedŽby Benjamin Franklin a U.S. diplomat in Paris at


.the time , who percipiently suggested it was causedŽ .by volcanism Franklin, 1784 . However, it is uncer-

tain if the Laki eruption was alone responsible forthis transient cooling because the coincidental erup-tion of the Asama volcano in Japan also in 1783;may have injected substantial amounts of SO into2

Ž .the stratosphere Sigurdsson, 1990b . The Laki erup-tion was minor in comparison with those found in

Ž .most CFBPs. Thus, Self et al. 1997 estimated thatindividual flows in the Columbia River Provincereleased roughly two orders of magnitude more SO2

than the Laki flow, although the 73 ka Toba eruptionprobably released a comparable amount of SO .2

However, unlike the Toba eruption, it is not clearthat fissure eruptions are capable of injecting gasesinto the stratosphere.

Fissure eruption is generally the most quiescentŽ .form of volcanic activity e.g. Self et al., 1997 ,

because gases are able to escape relatively easilyfrom the fluid lava. Thus, many workers doubt sucheruptions are capable of injecting gas into the strato-

Ž .sphere Pinto et al., 1989; Sigurdsson, 1990a,b .Only the presence of fire fountains along the lengthof fissures provides a potential mechanism of inject-ing gases higher into the atmosphere, particularlywhen they occur at higher latitudes where the tropo-

Ž .sphere is at its thinnest. Woods 1993 has postulatedthat convective plumes may rise above fire fountains

Žand entrain unsaturated water vapour assuming a.moisture-rich troposphere during eruption . Initially,

the thermal energy of the plume will cause it toascend, but as atmospheric pressure decreases a satu-ration height will be reached whereupon the watervapour will condense. The resultant release of latentheat will add further impetus to the plumes ascent.Such effects are however proportionately more im-portant for smaller fissure eruptions because atmo-spheric vapour is confined to the lower few kilome-tres of the atmosphere. Larger fissure eruptions maybe able inject volcanic gases into the stratosphere

Ž .due to the at least hypothetical possibility that theheat released from the surface of large lava flows

Žcould create large atmospheric convection cells De-.vine et al., 1984 . Such cells could generate hurri-

cane-strength winds at the surface of the flow. Suchwinds were not reported during the Laki eruptions,though it is possible that the larger flows of CFBPsmay have generated such circulation.

In summary, the eruption of volcanic gases hasbeen postulated to cause a succession of climaticchanges on a timescale of months to thousands of

Ž .years Fig. 1 . The short-term effects have beenobserved to occur following historical eruptions,whereas the longer term climatic impact of volcan-ism has not been clearly demonstrated. In essence,volcanic eruptions affect the weather, but not neces-sarily the climate. However, our modern perspectivedoes not take into account the voluminous fissureeruptions of ancient CFBPs which exceed, by manyorders of magnitude, the volcanic activity of histori-cal times. It is these huge eruptions that are widelyregarded as a cause of many extinction events. Tworecent advances in our knowledge of LIPs has con-siderably strengthened this connection, and both aredue to improvements in radiometric dating tech-niques. Firstly, many LIPs have been shown to record

Ž .brief bursts of volcanicity ;1 Ma , thus focussingtheir damaging effects in geologically short intervalsand, secondly, these intervals have been shown toclosely coincide with mass extinctions, as first noted

Ž .by Morgan 1986 .Before examining the relationship between indi-

vidual mass extinction events and volcanicity, it isfirst important to review the current debate on thegenesis of large igneous provinces. Several compet-ing models are available and their predictions sug-gest distinctly different environmental consequences.

3. The genesis of large igneous provinces

Many LIPs consist of subaerial sheet flows, mostcommonly composed of quartz tholeiites, whichrange in volume from several hundred to several

Žthousand cubic kilometres Baksi, 1990; Arndt et al.,1993; Coffin and Eldholm, 1993, 1994; Courtillot et

.al., 1999 . Total lava volumes for continentalprovinces are enormous and range from 1.7=105

km3 for the Columbia River Basalts to provinces thatmay originally have exceeded 2=106 km3. ManyCFBPs were erupted in geologically brief intervalsŽ .1–2 Ma although some, notably the Siberian Trapsand Brito–Arctic Province, were emplaced in two ormore distinct phases separated by quiescent intervals.The eruption history of the Columbia River Basalts

Ž .is well known Fig. 3 , and may be typical for many


Ž .Fig. 3. Eruption history of the Columbia River flood basalts from data in Baksi, 1990 . This small province is relatively well known andmay provide a model for the eruption histories of lesser known but much larger provinces.

provinces. In this region eruptions began at 17.2 Maand rapidly reached a peak lava accumulation ratearound 16.1 Ma before declining to a prolonged

Ž .AtrickleB Baksi, 1990 . The duration of the peakeruptive interval has been estimated to be as short as

Ž10,000 years for some provinces e.g. Officer et al.,.1987 , and it is this phase that is believed to cause

mass extinction.CFBPs, and other large igneous provinces, are

widely regarded to be the product of mantle plumesthat have ascended from deep within the mantle,impinged on the base of the lithosphere, and subse-

Žquently erupted e.g. Morgan, 1972; Loper and Mc-Cartney, 1986; Courtillot and Besse, 1987; Richardset al., 1989; White and McKenzie, 1989; Hill, 1991;

.Coffin and Eldholm, 1994 . Melt is produced duringdecompressive rifting and thus melt volumes are attheir greatest where the lithosphere is thin and able

Žto extend freely as at oceanic sites White, 1993;.Farnetani and Richards, 1994 . Melt generation mayŽalso be enhanced if eclogite from subducted oceanic

.crust is entrained in plumes during their ascentŽ .Cordery et al., 1997 . The initial arrival of theplume at the base of the lithosphere is postulated tocause widespread doming with elevations suggested

Žto range from 1–4 km Cox, 1989; Campbell andGriffiths, 1990; Hill, 1991; Farnetani and Richards,

.1994; White and Lovell, 1997 . This effect has beenlinked to the widespread sea-level fall that occurs

Žduring several mass extinction events e.g. Hallam,.1987; Officer et al., 1987; Erwin, 1993 . Conversely,

where plumes arrive at the base of the oceanicŽlithosphere the doming and the subsequent oceanic

.plateau formation will tend to displace water ontoŽcontinental shelves Larson, 1991a,b; Kerr, 1998;

.Hallam, 1999 .Not all LIPs conform to this plume-generated

scenario. Most significantly for mass extinction stud-ies, evidence for pre-eruption doming is often scant,

Žnotably for the Siberian Trap province Cordery et.al., 1997; Sheth, 1999 . Don Anderson and col-

leagues have even challenged the notion that mantleplumes are responsible for flood basalt provincessuggesting alternatively that they may be rapidlygenerated from a volatile-rich upper mantle at sites

Žthat have a prolonged history of prior rifting Ander-son et al., 1992; Anderson, 1994; King and Ander-

.son, 1995 . The competing models for CFBP genera-tion imply differing volumes of volatile generationŽe.g. the dry melting of the White and McKenzieŽ .1989 model contrasted with the volatile-rich magma

.source in Anderson’s model and differing conse-quences for sea-level changes: both factors are ofimportance in mass extinction studies.


Most controversially, it has been proposed thatLIPs may be generated at sites of bolide impactŽRampino, 1987; Alt et al., 1988; Stothers and

.Rampino, 1990 , thus uniting the two competingmodels for the end-Cretaceous mass extinction.However, mapping of the subcrop beneath the Dec-can Traps reveals no evidence for a crater structureŽ .Mitchell and Widdowson, 1991 while the Chicxu-lub impact post-dated the onset of eruptions in IndiaŽ .Bhandari et al., 1995; Bajpai and Prasad, 2000 .More generally, impact excavation is considered in-capable of generating the vast volumes of lava found

Žin flood basalt provinces White, 1989; Loper and.McCartney, 1990 .

ŽThe link between CFBPs and some other large.igneous provinces and mass extinction is examined

in chronological order and in a case-by-case basisbelow.

4. The Panjal Volcanics and Emeishan floodbasalts

The Panjal Volcanics are the oldest, and also themost deformed, of known Phanerozoic CFBPs. Theycover approximately 12 000 km2 of NW India andreach a peak thickness of 2.5 km in the Kashmir

Ž .Valley Honnegger et al., 1982 . Basaltic lavas andlesser volumes of rhyodacitic tuffs marked the mainphase of volcanism although minor alkali volcanicspersisted in the Kashmir region into the Early Trias-

Ž .sic Veevers and Tewari, 1995 . The basalt lavasoverlie Lower to Middle Permian subaerial tomarginal marine strata and they are in turn overlain

Ž .by early Late Permian Kazanian pelagic strataŽ .Nakazawa et al., 1975 . This facies change suggeststheir emplacement is probably related to rifting and

Žthe development of a distal passive margin Papritz.and Rey, 1989 . The subsequent deformation of the

Panjal Volcanics, and metamorphism to amphibolitegrade, makes estimation of the original volume ofthe province problematic although the region is un-likely to have constituted a Asuper provinceB of asize comparable with the Deccan Traps.

Contemporaneous and comparable volcanicity alsooccurred in South China with the eruption of theEmeishan flood basalts. These too are associated

Žwith a probable rifting event probably the Qiang-

.tang Terrane from the South China Block , whichbegan in the early Permian, and they have also beensubsequently deformed during the closure of TethysŽ . 5Chung et al., 1998 . Their outcrop covers 3.3=10km2 of SW China and consists of flood basaltstogether with lesser volumes of picritic flows and

Žtrachytic and rhyolitic tuffs Chung et al., 1998; Jin.and Shang, 2000 . The lavas average 705 m thick

and locally they reach 5 km thickness. Yin et al.Ž .1992 estimated the original volume of the province

6 3 Ž .to be 0.6=10 km , while Jin and Shang 2000suggested roughly half this figure.

The precise age of the Emeishan flood basaltsprovince has yet to be determined. The lavas rest on

Ž .Middle Permian carbonates the Maokou Formationand they are directly overlain by Late Permian and

Žlocally Early Triassic marine strata Yin et al., 1992;.Chung and Jahn, 1995 . Recently, Jin and Shang

Ž .2000 have reported the presence of fusulinidforaminifera from shallow marine limestones in-terbedded with basal lava flows, that unequivocally

Ž .indicate a Guadalupian Middle Permian age . Mag-Žnetostratigraphic results from a thin 532 m thick-

.ness development of the flood basalts in westernGuizhou Province indicates the lower 450 m waserupted during a single, normal polarity chron and

Žthe upper 82 m during a reversal Huang and Opdyke,.1996 . Thus, the Emeishan CFBP, like the Panjal

Volcanics, appear to have been erupted during aŽbrief interval Huang and Opdyke suggest less than 1.million years around the MiddlerLate Permian

boundary. However, acidic ash bands are widespreadin P–Tr boundary sections in South China, and theyare widely regarded to have been derived from the

ŽEmeishan Province Yin et al., 1992; Chung and.Jahn, 1995; Chung et al., 1998 , which implies a

prolonged interval of acidic volcanicity spanning theentire Late Permian and lasting into the Early Trias-

Ž .sic. However, thick 15–25 m , coarser-grained,acidic-intermediate tuff bands are also known from

Ž .the boundary interval in south Primorie SE Siberiasuggesting this area lay closer to the eruptive centrethan South China, where the tuffs are finer-grained

Ž .and an order of magnitude thinner Kozur, 1998 .Clearly, more work is needed on the provenance ofthese P–Tr boundary tuffs.

Attempts to relate the Emeishan flood basalts inŽthe end-Permian mass extinction e.g. Yin et al.,


.1992 are probably inappropriate, but the dating fromthe region suggests it may be better implicated, alongwith the Panjal Volcanics, with the end-Guadalupian

Ž .mass extinction Hallam and Wignall, 1997 . Thisearlier event, which has only been identified withinthe last few years, primarily affected equatorial ma-rine invertebrate taxa, notably the fusulinid

Žforaminifera Leven, 1993; Jin et al., 1994; Stanley.and Yang, 1994 . The best data come from the

southwestern United States and South China wherethe extinctions appear related to major regional re-gressions which eliminated the shallow-marine car-bonate habitats that were the repository for muchMiddle Permian marine diversity. The Mid–LatePermian boundary may mark an all time low-point of

Žsea level in the Phanerozoic Hallam and Wignall,.1999 . Within South China itself, this sea-level fall

could reflect regional doming prior to the eruption ofŽ .the Emeishan flood basalts Chung et al., 1998 , but

globally some other mechanism must be invoked.

5. Siberian Traps

5.1. The Õolcanics

The immense size of the Siberian Trap provincehas become a part of geological folklore and they are

Žoften depicted to cover vast areas e.g. Officer and.Page, 1996, p. 169 . In fact, flood basalts cover only

a modest 3.4=05 km2 of northwest Siberia, al-Ž .though pyroclastics and intrusives particularly sills

greatly increase the total area of the province to6 2 Ž1.5=10 km Zolotukhin and Al’Mukhamedov,

.1988; Fig. 4 . The original area of the flood basaltswas undoubtedly greater, although proposed figures

6 2 Žof 5=10 km Czamanske et al., 1998; Kozur,.1998 , are probably exaggerated. Estimates of their

original volume are even more speculative. Manyauthors assume 1–2 million km3, and Courtillot et al.Ž . 6 31999 suggest as much as 4=10 km , but only4=105 km3 currently remain. The discovery of athick pile of Early Triassic basalts in a borehole fromthe western Siberian depression suggests the provincemay extend westwards beneath the Jurassic–Creta-

Žceous sediments of this major basin Westphal et al.,.1998 . However, within the province itself, the dis-

tribution of lavas suggests that they do not constitutea single continuous province but rather the amalga-

Žmation of several AsubprovincesB Mitchell et al.,.1994 .

Volcanism in western Siberia began with the for-Žmation of the Tuffaceous Series Sadovnikov and

.Orlova, 1993, 1998; Fig. 5 . These predominantlybasaltic tuffs underlie the flood basalts in most of theprovince and in the south of the region they domi-

Ž .nate the entire succession Fig. 4 . The tuffs haveproved difficult to date, but the recent discovery of

Žconchostracan faunas within interbedded strata andtheir comparison with similar faunas in well-dated

.Chinese sections indicates they span the middle partŽ .of the Late Permian Kozur, 1998 . Flood basalt

flows overlie the tuffs over wide areas and recentradiometric dating efforts suggest the onset of erup-

Žtions was coincident with the Permian–Triassic P–. Ž .Tr boundary. Claoue-Long et al. 1991 obtained´

U–Pb dates of zircons from an ash band immediatelybelow the P–Tr boundary in South China of 251.2"

Ž .3.4 Ma. Renne et al. 1995 analysed sanidine fromthe same ash band and obtained an 40Ar–39Ar age of249.91"1.52 Ma, including external error. Redating

Ž .the same ash band again, Bowring et al. 1998obtained a U–Pb age of 251.4"0.3 Ma. These datessuggest the P–Tr boundary occurred between 250and 251 Ma, with a greater likelihood for the olderend of this age range. In contrast, initial 40Ar–39Ardates obtained from the Siberian Traps were substan-

Ž .tially younger Baksi and Farrar, 1991 , althoughŽ . ŽRenne and Basu’s 1991 248.4"0.3 Ma internal

. 40 39error only Ar– Ar date from a basalt flow nearthe base of the lava pile at Noril’sk was the first toindicate that the onset of eruptions was close in age

Žto the P–Tr boundary. Redating of the standard Fish.Canyon sanidine , used to calculate this age, pro-

duced an age of 250.0"1.6 Ma and therefore anŽ .even closer correspondence Renne et al., 1995 ,

including all uncertainties this date becomes 250.0"Ž .2.3 Ma Renne et al., 1998 . Within error this is

comparable to a U–Pb age of 251.2"0.3 Ma thatŽ .Kamo et al. 1996 obtained from an intrusion in the

lower third of the Noril’sk lava pile. Venkatesan etŽ .al. 1997 analysed samples from throughout the

Noril’sk succession and obtained ages of 247.1"1.9Ma for the base and 247.6"2.5 Ma for the top ofthe lava pile indicating a rapid eruption rate of


Ž .Fig. 4. Geological map of northwestern Siberia showing extent of Siberian Trap volcanics after Zolotukhin and Al’Mukhamedov, 1988 .

perhaps a million years or less. Venkatesan et al.Ž .1997 used a different standard to calculate their40Ar–39Ar ages, which cannot therefore be directlycompared with Renne et al.’s date, but it is clear thatfor the Noril’sk succession at least, the eruptionsbegan around the P–Tr boundary and persisted forapproximately a million years or less. This conclu-sion is essentially in accord with the available bios-tratigraphc dating from the region, which suggeststhat the onset of fissure eruptions occurred in the

Žlater part of the Dorashamian Stage the last stage ofthe Permian, which is also known as the Changxin-

.gian Stage , slightly below the palynologically de-Žfined P–Tr boundary Sadovnikov and Orlova, 1993,

.1998 . Conchostracan evidence suggests a similarŽ .age assignment Kozur, 1998 . This subtly older age

is below the resolution available from radiometricdating.

The timing of eruptions in the Noril’sk region isreasonably well constrained, but it is important toappreciate that this area only constitutes 7% of the

Žtotal volume of the province Venkatesan et al.,.1997 , and may not be representative of the entire

province. In the Maimecha–Kotui area in the NE ofthe province, up to 3 km of lavas infill a pre-existing

Ž .Late Permian graben Basu et al., 1995 . Unlike theremainder of the flows in the province, which aretholeiitic basalts, the Maimecha–Kotui volcanics


Fig. 5. Comparison of the timing of Siberian volcanic events in the Permian–Triassic interval with the contemporaneous carbon isotopeŽ . Ž .curve cf. Holser et al., 1989; Magaritz et al., 1992 . The alternative Triassic age of the main phase of Siberian Traps volcanism is after

Ž . Ž . Ž . Ž .Venkatesan et al. 1997 . The asterisks denote 1 the main level of marine invertebrate extinctions cf. Hallam and Wignall, 1997 and 2Ž .the level of freshwater invertebrate extinctions Kozur, 1998 . The correspondence in timing between the extinctions and flood basalt

volcanism is very close, although the best available radiometric ages suggest the main phase of volcanism slightly post-dates theend-Permian mass extinction.

consist of a chemically distinct suite of alkai-ultra-basics which appear to be older than the Noril’sk

Ž .succession. Thus, Basu et al. 1995 obtained an40Ar–39Ar plateau age of 253.0"2.6 Ma from anolivine nephelinite in the region. This date wascalibrated using the same Fish Canyon sanidine stan-

Ž .dard and the recalibrated age that Renne et al. 1995used to recalculate the ages they reported in Renne

Ž .and Basu 1991 . Thus, the age for the onset ofNoril’sk fissure eruptions and Maimecha–Kotuieruptions can be compared; they indicate that erup-tions in the latter region occurred 0.7–5.0 Ma earlier

Ž .at the 95% confidence level Venkatesan et al. 1997 .Clearly, much additional radiometric dating is re-quired from elsewhere in the province.

Much of our knowledge of the nature of SiberianTraps volcanism also derives from Noril’sk. Individ-ual flows in this area are small by the standards ofCFBPs and rarely exceed a few tens of metresthickness and a few tens of kilometres in extentŽ .Sharma, 1997 . Thus, the eruption history consisted

Žof numerous, small volume flows typically one flow3 .of 10 000 m every 1000 years . The Noril’sk lava

Ž .pile also contains an unusually high proportion 10%of basaltic pyroclastics. There are approximately 30beds in total, ranging from a few tens of centimetres

Ž .to more than 100 m thick Venkatesan et al., 1997 .Tuffs are also an important component of theMaimecha–Kotui volcanics, and in the succession to

Ž .the south of Putorana Fig. 4 .


The abundance of tuffs and the relatively smallvolume of individual lava flows are an unusualaspect of the Siberian Trap province. Also notewor-thy is their intracratonic setting. All other provinces,with the exception of the small Columbia RiverProvince, occur at locations marked by subsequent

Žcontinental rifting or former sites in the case of the.Emeishan and Panjal Volcanics . As noted above,

many authors ascribe this relationship to a plume-re-Žlated mechanism of continental break-up e.g. Cour-

.tillot et al., 1999 . However, some of the predictionsof the plume model are not found in the SiberianTraps, notably the absence of pre-eruption domingŽ .Czamanske et al., 1998 . Modelling of plumes ofthe size required to generate the Siberian Trapssuggests that up to 4 km of uplift should mark thearrival of the plume at the base of the lithosphereŽ .Farnetani and Richards, 1994 . There is no evidencefor such an event in western Siberia and in both theNoril’sk and Maimecha–Kotui areas the lavas infill apre-existing graben topography indicating that rifting

Žnot uplift predated the eruptions Zorin and.Vladimirov, 1989; Czamanske et al., 1998 .

Nonetheless, the plume model remains the preferredalternative in most studies of the Siberian TrapsŽRenne and Basu, 1991; Veevers et al., 1994;Conaghan et al., 1994; Coffin and Eldholm, 1994;

.Sharma, 1997 .

5.2. Extinction mechanisms

Whatever the ultimate origin of the Siberian Traps,their significance is well established due to the cen-tral role they play in the majority of current modelsfor the end-Permian mass extinction. The higher

Ž .proportion of tuffs relative to many other CFBPs inthe Siberian Traps have focussed attention on thelikely cooling effects of volcanic dust and sulphate

Žaerosols Campbell et al., 1992; Conaghan et al.,1994; Renne et al., 1995; Kamo et al., 1996; Kozur,

. Ž .1998 . Thus, Campbell et al. 1992 proposed thatcooling was sufficient to cause an intense glaciationfor the duration of the Siberian Traps eruption—aperiod of 600 ka in their estimation. This, in turn, issuggested to have been the cause of the widelyreported, major eustatic sea-level fall at the end of

Ž .the Permian e.g. Holser and Magaritz, 1987 . Alter-

Ž .natively, Erwin 1993 suggested that the fall mayrecord broad uplift, centred on western Siberia, im-mediately prior to the onset of volcanicity in theregion. However, as noted above, the lack of evi-

Ždence for uplift in the region Kamo et al., 1996;.Czamanske et al., 1998 makes this alternative un-

tenable. There is an equal lack of evidence for endŽ .Permian glaciation, although Campbell et al. 1992

suggested it may have been too brief to have leftphysical evidence. Recent reassessment of latest Per-mian and earliest Triassic sea-level changes suggeststhat the absence of such evidence may not be so

Ž .puzzling because rapid sea-level rise not fall isŽseen in most sections at this time Hallam and

.Wignall, 1999 . Hitherto, much of the evidence forsea-level fall in the latest Permian has been based onan absence of diagnostic latest Permian biostrati-graphic markers, but sequence stratigraphic analysisreveals the latest Permian–earliest Triassic interval

Žwas marked by a phase of rapid coastal onlap Haq.et al., 1987; Wignall and Hallam, 1993 . Hallam

Ž .1999 has speculated that the sea-level rise mayreflect a major pulse of intraoceanic volcanism. Thesame effect has been proposed for some Cretaceous

Ž .sea-level changes see below , but the absence of anypre-Jurassic oceanic crust makes this proposition dif-ficult to judge for the P–Tr sea-level changes.

If glaciation can be discounted as a factor in theend Permian extinction, the same cannot be said of

Ž .global cooling. Kozur 1998 in particular favoursthis mechanism as the main cause of the mass extinc-tion. His principal evidence is derived from thechanging geographic distribution of conodont speciesduring the crisis. In many Tethyan P–Tr boundarysections, the extinction interval is marked by theappearance of Clarkina carinata, a species inter-

Ž .preted to have cool water preferences Kozur, 1998 .However, the temperature preferences of Clarkinaspecies are by no means clear and it is noteworthythat other Clarkina species, such as the closelyrelated C. subcarinata, are inferred to have preferred

Ž .warm waters Kozur, 1998 . An alternative possibil-ity, that C. carinata was a deep water benthic formthat became more widespread during the deepeningcaused by the end-Permian sea-level rise, needs to beconsidered. In contrast to the conodont evidence,contemporaneous changes in the terrestrial flora sug-gest that cold-adapted forms faired particularly badly


during the end-Permian crisis. The eradication of theGlossopteris flora in high southern palaeolatitudes is

Ža particularly noteworthy feature of the event Retal-.lack, 1995 . Extinctions amongst Boreal marine in-

vertebrates in Spitsbergen were also at least as se-vere as those seen in the lower latitude sections of

Ž .Tethys Wignall et al., 1998 .The other climatic consequence of SO eruptions,2

acid rain, has also been postulated as a cause of oneof the more intriguing aspects of the extinction inter-val: the proliferation of fungi. Palynological samplesfrom P–Tr sections around the world reveal an in-crease of fungal spores in latest Permian times,particularly in low palaeolatitudes where they often

Ž .dominate assemblages Eshet, 1992 . Visscher et al.Ž .1996 attribute this to degradation, by volcanogenicacid rain, of floral ecosystems with the result thatfungi proliferated on the ample decaying vegetation.

Ž .Hallam and Wignall 1997 proposed an alternativeorigin linked to the extinction of many insect orders

Žat the end of the Permian Labandeira and Sepkoski,.1993 . Insects today, and probably in the Permian,

destroy a large proportion of the terrestrial flora, butin their absence fungi may respond to the increasedavailability of decaying vegetation. The possibilitythat the fungal spores event records the proliferationof marine rather than terrestrial fungi is a further

Žfactor that requires consideration Wignall et al.,.1996; Kozur, 1998 , and one that further weakens

the potential link between the fungal spike and acidrain.

A major phase of global warming is the mostobvious climatic signal associated with the end-Per-mian mass extinction. The evidence includes themigration of calcarous algae to Boreal latitudesŽ .Wignall et al., 1998 , the preferential loss of high

Ž .latitude floras noted above Retallack, 1995 , and thedevelopment of palaeosols typical of latitudes of 508

Žor lower at palaeolatitudes as high as 808S Retal-.lack, 1999; Retallack and Krull, 1999 . Oxygen iso-

tope data suggest equatorial temperatures may haverisen as much as 68C at the P–Tr boundary, althoughsuch inferences have to assume that salinity fluctua-

Ž .tions can be disregarded Holser et al., 1989 . Stron-tium isotope ratios in unaltered marine carbonatesand phosphates also point to an increase in atmo-spheric CO around this time. Following a low-point2

at the end of the Middle Permian, 87Srr86Sr ratios

began to rise at an increasingly rapid rate and reacheda highpoint around the end of the Early TriassicŽ .Martin and Macdougall, 1995 . This trend mayreflect increased continental weathering in an in-

Žcreasingly CO -rich atmosphere Martin and Mac-2.dougall, 1995 andror the enhanced leaching of

Ž .volcanogenic acid rain Conaghan et al., 1994 . Thealternative, that the Sr isotope curve reflects theenhanced continental erosion during major sea-level

Ž .fall Holser and Magaritz, 1987 , does not accountfor the lack of correlation between the eustatic sea-

Žlevel curve and the isotopic trends Hallam and.Wignall, 1997 .

Global warming may also have been the cause ofthe rapid development of marine anoxia seen in

Žmany latest Permian shelf sections Wignall and.Twitchett, 1996 . This AsuperanoxicB event is first

developed in the deep-water, pelagic chert sectionsfrom the accreted terranes of Japan at the end of the

Ž .Middle Permian Isozaki, 1997 , but it is the expan-sion of oxygen-poor conditions into shallow watersin the latest Permian which coincides with the ma-

Žrine mass extinction event Wignall and Hallam,.1992, 1996; Wignall et al., 1995 . Two effects of

warming, the decline of the equator-to-pole tempera-ture gradient and consequent decrease in oceaniccirculation, together with the lower solubility ofoxygen in warmer waters, may have been responsi-

Ž .ble for marine anoxia Wignall and Twitchett, 1996 .However, the ultimate cause of the elevated atmo-spheric CO concentrations required to generate the2

warming is not known. Volcanic CO derived from2

the Siberian Traps is the preferred source of manyŽe.g. Campbell et al., 1992; Conaghan et al., 1994;

.Veevers et al., 1994; Wignall and Twitchett, 1996 ,Žalthough contributions from gas hydrates Erwin,

.1993; Morante, 1996 , and the oxidation of Gond-Ž .wanan coals Faure et al., 1995 , may also have been

important.

5.3. Carbon isotopic trends

Dramatic d13C fluctuations in both marine and

terrestrial sections during the P–Tr interval may holda clue to the cause of atmospheric CO changes.2

Values of d13C in marine carbonates decreased grad-

ually through the Dorashamian Stage before a rapid


negative excursion, of 4–5‰ magnitude, began im-Žmediately below the P–Tr boundary Holser et al.,

1989, 1991; Xu and Yan, 1993; Morante et al., 1994;.Morante, 1996 . This negative spike indicates a brief

influx of isotopically light C into the ocean-atmo-sphere system around the time of the end-Permianmass extinction. Volcanogenic CO is a potential2

13 Žsource, although its d C value of y5‰ McLean,.1985 is probably not sufficiently light to achieve the

observed swing. This is readily appreciated by com-paring the amount of C in surficial carbon reservoirs,

19 Ž .estimated to be around 5=10 g C Berner, 1999 ,with the estimated volumes of CO emitted during2

the formation of CFBPs. Based on measurementsŽ .from Hawaiian eruptions, McCartney et al. 1990

calculated that 1 km3 of basalt emits 5=1012 g ofC. Thus, if a high value of 2=106 km3 of basalt isassumed to have been originally present in theSiberian Traps, then 1=1019 g of carbon dioxide Cmay have been released during their eruption. This is

sufficient to only cause roughly 20% of the observedisotopic swing in the latest Permian. There are clearlynumerous assumptions implicit in this reasoning, notleast the y5‰ isotopic composition of volcanicCO . In this context, the debate on the ultimate2

origin of CFBPs is apposite to this assumption.Ž .Anderson 1994, 1999 suggested that, rather than

being sourced from deep mantle plumes, CFBPs mayresult from the melting of a shallow mantle enrichedin recycled lithosphere. If this layer, termed theperisphere, is enriched in isotopically light, sub-ducted organic C then the d

13C of the volcanic CO2

emitted from this source may be considerably lighterthan y5‰ value. Eruption of around 1=1019 g Cwith a d

13C value of y20‰ would be sufficient tocause the isotope excursion seen at the end of thePermian. However, before accepting these specula-tions, it is salient to recall that much, if not most, ofthe Siberian Traps province was erupted in the EarlyTriassic after the d

13C excursion.

Fig. 6. Chain of events caused by the eruption of the Siberian Traps which could serve as a model for all volcanogenic extinctions. Note thatŽthe postulated cause-and-effect links shown on the left-hand side of the diagram are not supported by geological evidence perhaps because

.they were of too brief duration . The global warming effects are supported by geological evidence although the magnitude of C isotope shiftsuggests that other sources of CO , such as gas hydrates, must be implicated. Thus, the ability of volcanic CO to trigger further CO2 2 2

release may have been the single most important cause of climatic change during the P–Tr events, and perhaps during other intervals of LIPformation.


The sudden influx of light C into the oceans andatmosphere of the latest Permian has also been at-tributed to a near-total collapse of primary productiv-

Žity Holser and Magaritz, 1992; Magaritz et al.,.1992; Wang et al., 1994 . However, consideration of

the masses involved once again highlights that suchan event is unlikely to cause more than a smallnegative shift. Thus, the modern biomass amounts to8.3=1017 g C and has an isotopic composition ofaround y20‰ which, if it were to be added to the5=1019 g C present in the inorganic C componentof the ocean–atmosphere system, would clearly havelittle impact on d

13C values.In recent years, it has been appreciated that one of

the major repositories of isotopically light C occursin the form of highly volatile methane hydratesŽ .clathrates buried at shallow depths beneath cold

Ž .andror deep seas Dickens et al., 1997 . HydratesŽ 13 .are isotopically very light d Csy65‰ , and there

may be as much as 1=1019 g of C present inhydrates beneath modern seas. If only 10% of thismaterial were to be released to the atmosphere asmethane, it would be sufficient to cause the d

13CŽshift observed across the P–Tr boundary Erwin,

.1993; Bowring et al., 1998 .Clearly, the magnitude of the end-Permian d

13Cexcursion renders its origin somewhat problematic.Eruption of vast volumes of volcanic CO and pro-2

ductivity shutdown may have occurred at this time,but these events can only account for a fraction ofthe observed C isotope changes. However, the warm-ing effect of the CO release may have triggered the2

dissociation of massive amounts of methane hy-drates, and thereby produced the d

13C excursion andŽ .exacerbated the warming trend Fig. 6 . As noted

above, there is substantial geological and palaeonto-logical evidence for globally warm conditions in theEarly Triassic. Similar Arunaway greenhouseB sce-narios have also been proposed for the Cenoman-ian–Turonian interval of the Cretaceous and in the

Ž .late Palaeocene see below with flood basalt volcan-ism implicated in both cases.

6. Central Atlantic Magmatic Province

The break-up of the central Atlantic region in theearly Jurassic was marked by extensive volcanism,

of which the flood basalts of the Newark Basin, inthe NE of the United States, are the best knownŽ .Olsen et al., 1997 . The original volume of theseextrusives was probably a relatively modest 50 000

3 Ž .km McHone, 1996 , although study of scatteredoutcrops in French Guyana, Surinam and Guinea led

Ž .Deckart et al. 1997 to speculate that a major CFBPmay have formed prior to central Atlantic rifting.This hypothesis has received further support from

Ž .the work of Marzoli et al. 1999 , who discoveredadditional contemporaneous flood basalts in northern

Ž .and central Brazil Fig. 7 . A Central Atlantic Mag-Žmatic Province has therefore been proposed Marzoli

.et al., 1999 , with scattered outcrops now cen-tripetally located around the periphery of the central

Ž .Atlantic region Fig. 7 . Estimates of the originalarea of the province, based on the assumption thatthe flood basalts originally formed a continuous en-tity, range up to 7=106 km2, with a volume of at

6 3 Ž .least 2=10 km Marzoli et al., 1999 . If this isthe case, then the province may originally have been

Ž .one of the largest of its kind Olsen, 1999 .The best constrained dates for the Central Atlantic

Magmatic Province are from the Newark Basin whereseveral U–Pb ages have suggested the onset of vol-

Žcanism began at 201"1 Ma Dunning and Hodych,.1990; Weems and Olsen, 1997 and lasted only

Ž .580"100 ka Olsen et al., 1997 . Extensive dykesin southeastern USA yielded 40Ar–39Ar ages, with

Žexternal errors, of 199.5"2.0 Ma Hames et al.,.2000 . Less well-dated lavas and intrusives of South

America suggest a similar age with U–Pb ages rang-Ž .ing from 204 to 195 Ma Deckart et al., 1997 and

Ž .201 to 197 Ma Marzoli et al., 1999 . The mostprecise dates place the eruptions within the Hettan-gian Stage of the basal Jurassic and this is con-firmed, in the Newark Basin, by the occurrence ofthe lowest lava 30 m above the base of the palyno-

Ž .logically defined Triassic–Jurassic T–J boundaryŽ .Fowell and Olsen, 1993 . The loss of 60% of paly-nospecies at this boundary, together with a prolifera-tion of fern spores, suggests a sudden and severecrisis in the terrestrial flora at the boundary. An even

Ž .more severe )95% species extinction loss hasbeen reported for leaf species in northern EuropeŽ .McElwain et al., 1999 . However, this floral extinc-tion has yet to be recognised beyond the North

Ž .Atlantic region Hallam and Wignall, 1997 .


ŽFig. 7. Global distribution of continental flood basalt provinces and oceanic plateaus cf. Coffin and Eldholm, 1994; Farnetani and Richards,.1994 .

Ž .The close but not perfect temporal link betweenthe onset of flood basalt eruptions and a terrestrialcrisis has led most workers to infer a cause-and-ef-

Žfect relationship Rampino and Stothers, 1988; Cour-.tillot et al., 1999; Olsen, 1999; Palfy et al., 2000 .´

Major climatic changes have been reported from theT–J boundary interval. Thus, McElwain et al.’sŽ .1999 analysis of stomatal density in fossil leavesfrom E. Greenland and the Baltic suggests a majorincrease of atmospheric CO in the basal Jurassic2

which may be attributable to central Atlantic volcan-ism.

Further intriguing data on the T–J mass extinctioncomes from high precision U–Pb dating of an ashlayer from a marine boundary section in the Queen

Ž .Charlotte Islands, Canada, where Palfy et al. 2000óbtained a date of 199.6"0.3 Ma from a level 5 mbeneath the top of the highest radiolarian zone of theTriassic. This appears to be somewhat younger thanthe terrestrial T–J boundary in the Newark Basin,

Ž .which led Palfy et al. 2000 to suggest that marine´

extinctions may have post-dated the terrestrial onesby approximately 700 ka. However, the precise dat-ing of the end-Triassic mass extinction in marine

Žsections is not well established Hallam and Wignall,.1997 . Bivalves and ammonoids are amongst the

most prominent victims of this crisis and their recordsuggests a gradual decline in the latest TriassicRhaetian Stage, which may be partly due to localfacies changes, with a final coup-de-grace somewhat

Žbelow the T–J boundary e.g. Hallam and Wignall,.2000 . If the marine T–J boundary is indeed

marginally younger than the terrestrial boundary,then the latest Rhaetian marine extinction probablycoincides with the end-Rhaetian floral crisis.

7. Karoo and Ferrar Traps

Early Jurassic break-up of Gondwana saw theemplacement of a major CFBP now divided by theSouth Atlantic, the Karoo Traps in South Africa andthe Ferrar Traps in Antarctica which together contain


in excess of 2.5=106 km3 of lava. Dating of theKaroo lavas indicates a brief eruption interval around

Ž .183"1 Ma Duncan et al., 1997 , the range in-creases to "2 Ma at 2s when external errors are

Ž .included Palfy and Smith, 2000 . A U–Pb age fromá marine ash layer dates the middle Toarcian as

Ž .181.4"1.2 Ma Palfy et al. 1997 , and Palfy and´ ´Ž .Smith 2000 have extrapolated an age of around 183

Ma for the base of the Falciferum Zone, an intervalassociated with significant extinctions.

The early Toarcian extinction was first identifiedŽ .in marine sections in NW Europe Hallam, 1961

and the subsequent discovery that extinctions alsoŽ .occur in South America Aberhan and Fursich, 1996¨

suggests that this is a global biotic crisis that particu-Žlarly affected shallow marine molluscs Little and

.Benton, 1995 . The widespread development of oxy-gen-poor conditions appears the most likely cause of

Ž .the crisis Hallam, 1987 . Once again the eruption ofvolcanic CO , and the consequent global warming,2

has been postulated as the ultimate origin of theŽAanoxicB event and thus the extinction itself Jenkyns,

.1999 . A 2–3‰ positive C isotope excursion hasbeen widely reported from the Falciferum ZoneŽ .Jenkyns, 1988 , which is generally regarded to re-flect the burial of isotopically light organic C in the

Ž .anoxic seas. Recently, Hesselbo et al. 2000 havereported a brief negative d

13C excursion of 2–3‰magnitude in the earliest Falciferum immediatelyprior to the positive excursion. As with the negativeexcursion at the P–Tr boundary, the magnitude ofthe early Falciferum event is too great to be ac-counted for by volcanic CO emissions even by a2

province the size of the Karoo–Ferrar Traps. Hes-Ž .selbo et al. 2000 speculate that the global warming

effect of volcanic CO may have triggered dissocia-2

tion of up to a quarter of the gas hydrate reservoir—Ž .the same scenario proposed by Bowring et al. 1998

Ž .for P–Tr events. Hesselbo et al.’s 2000 modelneatly incorporates several aspects of Toarcian geol-ogy. However, it is important to note that the highestresolution d

13C curve for the Falciferum Zone, ob-tained from belemnites, shows many more fluctua-

Žtions in this interval with four isotopic minima Mc-.Arthur et al. 2000 . The greatest excursion, and

Ž .probably the one identified by Hesselbo et al. 2000occurred in the mid-Falciferum Zone. If this is the

Ž .case, then the interval of volcanically triggered?

gas hydrate release occurred slightly after the onsetof oceanic anoxia and mass extinction.

8. Parana and Etendeka flood basalts´

The continued rifting of Gondwana in the EarlyCretaceous was associated with the eruption of amajor CFBP, now separated by the South Atlantic,the Parana flood basalts of South America and the´

Žsmaller Etendeka Traps of Namibia Harry and.Sawyer, 1992; Jerram et al., 1999 . The twin

provinces cover 1.5=106 km2 and may contain up6 3 Žto 2.35=10 km of extrusive volcanics Gladc-

.zenko et al., 1997 . Eruptions in the Parana Province´Ž .began at 133"1 Ma Renne et al., 1992 , and

40Ar–39Ar ages from throughout the province indicatethere was little or no diachroneity to the eruptions,thus the entire province is though to have formed in

Ž .0.6"1 Ma Renne et al., 1996 . These eruption agesfall within the Valanginian and Hauterivian StagesŽ .cf. Gradstein et al., 1994 , an interval associated

Ž .with low extinction rates Sepkoski, 1996 . Someworkers have suggested the eruptions may be linked

Žwith an end Jurassic mass extinction Rampino and.Stothers, 1988; Courtillot, 1994, 1999 . However,

this interval is considerably earlier and the so-calledŽ .extinction, first recognised in Sepkoski’s 1982

Žcompilation, is probably an artefact cf. Hallam and.Wignall, 1997 .

9. The Ontong Java Plateau

The Early Cretaceous also saw the emplacementof the largest single volcanic province on Earth, the

Ž .Ontong Java Plateau in the SW Pacific Fig. 7 ,which covers 2=106 km2 of the SW Pacific and

Žhas a crustal thickness of up to 35 km Saunders et.al., 1996 . Partial obduction has resulted in exposure

of sections of the plateau on the island of Malaita,these reveal basalt flows 20–70 m thick interbeddedwith limestones, therefore indicating submarine em-

Ž . 40 39placement Saunders et al., 1996 . Ar– Ar datesfrom the central part of the plateau suggest the upper

Žpart of the province formed around 122 Ma Maho-.ney et al., 1993 , indicating a late Barremian age in

Ž .the timescale of Gradstein et al. 1994 . This reason-ably accords with biostratigraphic evidence: plank-


tonic foraminifera from sediments at the top of thelava pile belong to the Globigerinelloides blowi Zone

Žof the succeeding Aptian Stage Tarduno et al.,.1991 . During the interval preceeding this zone, from

the late Barremian to the early Aptian, oxygen-poordeposition appears to have been widespread in the

Ž .world’s oceans Bralower et al., 1994 . This is theSelli Event, the first of several Cretaceous Aoceanicanoxic eventsB and, like the later events, it coincideswith a global sea-level rise. This eustatic change maywell be caused by the displacement of ocean waterscaused by the submarine emplacement of the Ontong

Ž .Java Plateau Tarduno et al., 1991 . The widespreaddevelopment of oxygen-deficient deposition has also

Žbeen indirectly linked with volcanism Keith, 1982;.Larson, 1991a,b; Jenkyns, 1999 . Volcanic CO re-2

lease and consequent global warming is the mostlikely cause of the anoxia, whereas the impact ofSO release is likely to have been negligible given2

that the eruptions were submarine.The early Aptian clearly has all the hallmarks of a

major environmental crisis and yet there is no associ-Ž .ated extinction event Hallam and Wignall, 1997 .

Once again the Early Cretaceous biota was immuneto the effects of a major volcanic episode. Minorextinctions, particularly of the reef-forming rudistbivalves occurred in the latter part of the AptianŽ .Hallam and Wignall, 1997 . These could coincidewith the formation of the Kergulean Plateau, anotheroceanic LIP. However, much further work needs tobe done on the dating of both the plateau and theextinctions to verify this coincidence.

10. The Caribbean–Colombian Plateau andMadagascar flood basalts

The most widespread phase of ocean anoxiardys-oxia in the Cretaceous occurred at the Cenomanian–

Ž . Ž .Turonian C–T boundary Wignall, 1994 . This wastemporally associated with a modest-sized extinctionevent, especially amongst deep-sea benthic

Ž .foraminifera Kaiho and Hasegawa, 1994 , althoughwhether there was a causal relationship is a matter of

Ždebate Banerjee and Boyajian, 1996; Hallam and.Wignall, 1997 .

Volcanism has long been implicated as a potentialŽcause of the C–T anoxic event Keith, 1982; Vogt,

.1989 , but the location of the volcanics has not, untilrecently, been certain. In fact several major volcanicprovinces may have formed during the C–T interval,

Ž 6namely the Caribbean–Colombian Plateau 4=103. Ž .km Kerr, 1998 , Broken Ridge in the Indian Ocean

Ž 6 3.2=10 km , perhaps a component of the OntongJava Plateau and a minor CFBP in MadagascarŽ . 6 3Storey et al., 1995 . A total of up to 20=10 kmof intraoceanic, plume-related basalts are thought tobe of this age. However, currently available radio-metric ages for these provinces suggests a selectionof emplacement times within the Turonian and there-fore slightly after the C–T events. Thus, the C–Tboundary occurs within the 93–90 Ma intervalŽ .Gradstein et al., 1994 while the Caribbean–Col-ombian Plateau ages average 89.5"0.3 Ma, theMadagascan CFBP dates between 88.5"2.9 and

Ž40 3987.6"2.9 Ma Ar– Ar dates with external error,Ž ..from Storey et al. 1995 , and the Late Cretaceous

portion of the Ontong Java Plateau formed around 90Ž .Ma Sinton and Duncan, 1997 . Despite this signifi-

Ž .cant mismatch in timing, Kerr 1998 has proposed aplausible scenario for volcanically driven globalwarming in the late Cenomanian. The submarineeruptions may have caused direct hydrothermalwarming of the oceans together with indirect warm-ing due to the release of large volumes of volcanicCO to the atmosphere. Further CO release would2 2

follow from acidification of the oceans due to vol-canic SO emissions at a rate that Kerr suggests may2

have been as high as 3=1017 kg yeary1. Warmingof the oceans would release further CO to the2

atmosphere with the end result being a ArunawaygreenhouseB for which there is ample palaeontologi-cal evidence, including the brief appearance of

Ž .crocodiles at the north pole Tarduno et al. 1998 .The reversal of this warming trend occurred abruptly

Ž .in the early Turonian Kuypers et al., 1999 , perhapsdue to a negative feedback in the carbon cycle in

Ž .which warmer and more humid conditions enhancenutrient input to the oceans, increase primary produc-

Žtivity and thus elevate organic C burial rates Jenkyns,.1999 . More directly, submarine volcanism may in-

crease iron availability in the oceans and thus stimu-Ž .late productivity Sinton and Duncan, 1997 .

In summary, Kerr’s model elegantly accounts formany aspects of C–T environmental change but,somewhat ironically, the eruption ages of the impli-


cated volcanic provinces suggests that they coincidewith the rapid cooling in the early Turonian and notthe extreme warming that preceeds this event.

11. Deccan Traps

11.1. The Õolcanics

The Deccan Traps cover 0.5=106 km2 of NWŽ .India Fig. 7 , and reach a peak thickness of 2.5 km

in the western outcrops of the Western Ghats regionŽMitchell and Widdowson, 1991; Venkatesan et al.,

.1993; Prasad and Khajuria, 1995 . Estimates of theiroriginal area range from 1.5–2.5=106 km2 andtheir original volume is generally estimated at 2=

6 3 Ž .10 km e.g. Widdowson et al., 1997 , althoughŽ .Officer et al. 1987 considered this to be an underes-

timate. Little known, but potentially extensive basaltflows extend offshore and may considerably increase

Žthe volume of the province Coffin and Eldholm,.1994 . The traps are divided into 13 formations and

mapping of their southern outcrops reveals south-Žward overstepping of flows Mitchell and Widdow-

.son, 1991 , a phenomenon that suggests a southwardmigrating eruption centre as India migrated north-wards over a stationary hotspot. Alternatively, it mayrecord progressively more extensive lava sheets dur-ing the eruption history. The former alternative im-plies a prolonged eruption interval, a conclusion ofconsiderable relevance to the debate on the cause ofthe K–T mass extinction.

Ž .The presence of well developed palaeosols bolesand lacustrine sediments within the Deccan Trapssuccession indicates substantial intervals betweeneruptions. The boles become more numerous towards

Ž .the top of the lava pile Widdowson et al., 1997suggesting that the eruption history may have beentypical of many CFBPs with a voluminous initial

Žburst of volcanism followed by a rapid decline cf..Fig. 3 . However, detailed examination of some sup-

posed boles reveals them to be altered pyroclasticflows indicating an under-appreciated component ofexplosive volcanism during Deccan Trap formationŽ .Widdowson et al., 1997 .

Both magnetostratigraphic and radiometric datingŽ .suggest a short period F1 Ma for the main erup-

tions. The majority of the lavas were erupted duringa single geomagnetic field reversal that Courtillot et

Ž .al. 1986 considered to be C29R, the 0.5-Ma-longŽ .chron that straddles the K–T boundary Fig. 8 . This

conclusion is supported by radiometric dating oflavas from Western Ghats that indicates a total erup-tion interval of less than 2 Ma sometime between 69

Žand 65 Ma Duncan and Pyle, 1988; Courtillot et al.,.1988 . Most subsequent studies have assumed a

roughly 1 Ma eruptive interval beginning at orslightly before the K–T boundary, but these ageassignments have not been without controversy.

Ž . 40 39Baksi and Farrar 1991 recalculated the Ar– ArŽ .ages of Courtillot et al. 1988 to suggest a consider-

ably longer eruptive phase from 67.6"1.8 to 64.5"0.5 Ma. Even more controversially, Venkatesan et

Ž .al. 1993 redated the Western Ghats sections andconcluded that the lower 1.8 km of lava were eruptedaround 67 Ma in Chron 31R: a date in the early lateMaastrichtian that is obviously considerably before

Ž .the K–T mass extinction Fig. 8 . Feraud and Cour-´Ž .tillot 1994 challenged these conclusions on the

grounds that Venkatesan et al.’s error bars wereinsufficiently large and therefore could not rule out aK–T boundary age. Recent very high precision Re–Os isochron dates indicate eruptions began at 65.6"0.3 Ma, thus confirming the coincidence with the

Ž .K–T boundary Allegre et al., 1999 . Nonetheless,`40Ar–39Ar dating of feeder dykes to the south of theDeccan Traps indicates volcanic activity persisted

Ž .until 62.8"0.2 Ma internal error only , a Danianage, indicating that 1 Ma is an unduly short estimate

Ž .of the eruptive interval Widdowson et al., 2000 ,although the main peak of eruption may still haveoccurred briefly at the K–T boundary.

11.2. Effects of Õolcanic gas emissions

Ž .McLean 1985 has estimated the volumes of CO2

released during Deccan volcanism using Leavitt’sŽ . Ž .1982 formula see above . Utilising an estimatedfigure of 2.6=106 km3 for the original volume ofthe Deccan Traps, McLean estimated that 5=1017

Ž 18 .m CO 6=10 g of C were released in 1.36 Ma.2

A similar figure is obtained if McCartney et al.’sŽ .1990 value, obtained from Hawaiian measure-

12 Ž .ments, of 5=10 g C as CO released per cubic2

kilometre of basalt, is used. The impact on theclimate of these gas volumes is difficult to predict.Not all the Deccan CO released would have re-2


Fig. 8. Comparison of geological and palaeontological events around the K–T boundary. Only the onset of a warming trend in the latestMaastrichtian appears linked to the Deccan Trap eruptions. Extinctions are best related to cooling in the earlier Maastrichtian and bolideimpact at the K–T boundary.

mained in the atmosphere during the eruption inter-val because feedback loops, particularly increasedweathering rates in a more CO -rich atmosphere,2

would draw down levels over a timescale of 10–100Ž .ka Caldeira and Rampino, 1990; Berner, 1999 .

However, other factors in the late Cretaceous worldmay have exacerbated the effects of these eruptions;for example, notably higher ocean temperatureswould have reduced their capacity to remove CO2

Ž .from the atmosphere McLean, 1985 . Caldeira andŽ .Rampino 1990 modelled the predicted mean global

temperature rise for Deccan Trap CO release using2

a range of starting conditions. In their worst casescenario, a 28C temperature rise lasting 0.5 Ma wasachieved using McLean’s figure for CO release, an2

eruption interval of only 10 ka and pre-eruptionatmospheric CO levels of 400 ppm. Assuming a2

more realistic 1 Ma eruptive phase, the temperatureincrease would be less than 18C over a 1-Ma inter-val, and the rise would be even less if a more

ŽCO -rich atmosphere is assumed Caldeira and2.Rampino 1990 . These are hardly the climatic

changes that one would expect to cause a mass

extinction and many workers have, perhaps not sur-prisingly, favoured volcanic SO as the principle2

cause of environmental deterioration.Basalts are amongst the most S rich of all lavas

and the Deccan Traps are likely to have injected18 Ž6=10 g of S into the atmosphere McCartney et. 21al., 1990 . In comparison with the 3.7=10 g of

sulphate in the oceans, this is a minor amount,although the acidification of freshwater systems mayhave been severe and a transient reduction of oceanicsurface water pH could have occurred. Officer et al.Ž .1987 calculated that surface water alkalinity mayhave been lowered by up to 10%, although thisfigure is based on the unrealistic assumptions that allthe volcanic gases were rained directly into the sea,rather than a combination of land and sea, and thatthe main phase of the eruption lasted only 10 ka.

Assuming the Deccan fissure eruptions were ca-Žpable of injecting gases into the stratosphere see

.above , short-term cooling could also have followedeach flow which may have triggered long-term cool-ing if the spacing of eruption events was sufficiently

Ž .close Cox, 1988 . As already noted, Officer et al.


Ž .1987 assumed a 10-ka peak eruptive interval intheir extinction mechanism. However, the presence

Žof intertrappean sediments Jaeger et al., 1989;.Venkatesan et al., 1993; Prasad and Khajuria, 1995

and, in the upper part of the succession, well devel-Ž .oped boles Widdowson et al., 1997 suggests this is

an unrealistically short interval. A more realisticmillion year peak eruptive interval implies an aver-age 1.7=1013g H SO ryear were produced in the2 4

atmosphere as a result of the Deccan Traps erup-tions. If the eruptions occurred as a series of lava

3 Žflows of up to 10 000 km Courtillot’s, 1990 figure. 17for Deccan flows , this implies a maximum of 10 g

of sulphate aerosols were injected into the atmo-sphere once every 10–100 ka over a period of 1

Žmillion years cf. Bhandari et al., 1995; Widdowson.et al., 1997 . Assuming that individual eruptions

Ž .lasted only 1–2 years, and rather unrealisticallythat all the SO was injected into the stratosphere,2

Žthen up to 18C global cooling may have occurred cf..Fig. 2 . For comparison, the bolide impact at Chicxu-

lub is thought to have injected at least 1018 g ofŽsulphate into the atmosphere Sigurdsson et al., 1992;

.Brett, 1992 and, unlike the uncertainty concerningfissure eruptions, there is no doubt that a bolideimpact would inject gases into the stratosphere.

11.3. The fossil eÕidence

It is unclear, from calculations of Deccan gasvolumes and their effects, whether the formation ofthis province can in any way be implicated with thecontemporaneous K–T mass extinction. However,potential links may also be found from investigationof the nature and timing of changes in the fossilrecord. Somewhat surprisingly, the fossil record fromthe Deccan Traps provides little evidence for anyvolcanogenic catastrophe. The principal environmen-tal change, recorded in intertrappean lacustrine sedi-ments, is the development of a climate of AmockaridityB due to the absence of vegetation cover on

Ž .fresh lava surfaces Khadkikar et al., 1999 . Thesame sediments also contain a freshwater fauna offish and amphibians which remain little changed

Žthroughout the Deccan Trap lava pile Jaeger et al.,.1989 . However, freshwater faunas were relatively

unaffected during the K–T mass extinction in NorthŽ .America as well Archibald, 1996 , and so are not a

good monitor of the event, although they suggest thatacid rain effects of volcanism were not significant.

Dinosaurs are the best known victims of theextinction and they too are known from India wherethey occur in the Lameta Group, beneath the Deccan

ŽTraps, and in the intertrappean sediments Jaeger et.al., 1989; Prasad and Khajuria, 1995 . There is a

slight decrease of dinosaur diversity above the baseof the Deccan Traps, perhaps due to the AmockaridityB in the region, followed by their rapid disap-pearance in the upper part of the succession. An Iranomaly occurs in intertrappean sediments from theupper part of the lava pile in Kutch Province but

Ž .dinosaurs specifically egg-shell fragments occursŽ .above this level Bajpai and Prasad, 2000 . Thus,

dinosaurs may have survived into the Tertiary inŽ .India, although Bajpai and Prasad 2000 alterna-

tively proposed that the Deccan Ir anomaly may beof volcanic origin and of latest Cretaceous age.

Comparison of the timing of events from else-where in the world also fails to provide a close link

Žbetween Deccan Trap eruptions and extinctions Fig..8 . Major climatic changes in the latest Cretaceous

began with a rapid cooling in the mid-Maastrichtianthat may have caused the extinction of several lowlatitude groups, notably the rudist bivalves and many

Žbenthic foraminifera MacLeod and Huber, 1996;.Abramovich et al., 1998; Fig. 8 . The diverse and

abundant inoceramid bivalves also went extinct dur-ing this interval, perhaps as a result of the same

Ž .cooling event Barrera, 1994 , although their demisein the Globotruncana gansseri foraminifer zone

Ž .slightly predates cooling Marshall and Ward, 1996 .The mid-Maastrichtian extinctions and cooling pre-date the onset of Deccan Traps eruptions by 4–6 MaŽunless the interpreted ages of Venkatesan et al.Ž . .1993 are correct , and are therefore unlikely to berelated. Changes in oceanic circulation, following thebreaching of tectonic sills in the South Atlantic andthe establishment of intermediate and deep water

Žflow, appears a likely cause Frank and Arthur,.1999 . The cooling trend persisted into the

Palaeocene but was punctuated by a 0.5 Ma warmingŽtrend in the latest Maastrichtian Barrera and Huber,

.1990 , that saw the expansion of low latitude plank-Žtonic foraminifera into mid-palaeolatitudes Pardo et

.al., 1999 . This interval is also noteworthy for amajor fall and then rise of eustatic sea level in the


last 100 ka of the Maastrichtian; the low-point oc-curred perhaps 10 ka before the K–T boundary andsea level was just beginning to rise at the boundaryŽHaq et al., 1987; Hallam, 1987; Hallam and Wig-

.nall, 1999 . This period of warming coincides withthe early phase of Deccan Trap eruptions and sug-gests that it may have been triggered by volcanicCO release.2

We have now arrived at the crucial mass extinc-tion interval for which there is a plethora of data anddebate. Many ammonites appear to have gone extinct

Žduring the eustatic lowstand Marshall and Ward,.1996 , but the most spectacular events were the

near-total extinction of planktonic foraminifera, at alevel marked by the all-important Ir anomaly, and

Žthe collapse of marine primary productivity Hsu and¨.McKenzie, 1985; Holser and Magaritz, 1992 . The

increased flux of volcanic CO and SO into oceanic2 2

surface waters may have caused an acidity increasesufficient to eliminate planktonic groups, particularly

Žthe pH-sensitive planktonic foraminifera Officer et.al., 1987; McCartney et al., 1990 . However, details

of the timing of this event suggests that bolideimpact was more likely the cause. Amongst theever-more-detailed studies of the boundary interval,

Ž .Kaiho et al. 1999 have provided a valuable studyfrom the Caravaca section in Spain. The extinctionof planktonic foraminifera coincides with a declineof the surface-to-deep d

13C gradient which is widelyheld to signify the near-elimination of primary pro-

Žductivity in the oceans Hsu and McKenzie, 1985;¨.Holser and Magaritz, 1992 . In Caravaca this isotope

excursion is confined to the 5 mm of sediment abovethe sharply defined Ir anomaly. Based on the average

Ž .sediment rates for this section, Kaiho et al. 1999inferred a 13-ka interval of productivity shutdown.Oxygen isotope ratios also change dramatically im-mediately above the Ir anomaly with up to 58Cwarming of surface waters indicated. This eventpersisted a few thousand years longer than the d

13Cexcursion and may indicate that the CO -driven2

warming was only reversed once oceanic productiv-ity had been reestablished, thereby allowing signifi-cant removal of C to the seafloor sediments.

Detailed sampling of many other K–T boundarysediments reveals comparable, and equally rapid en-vironmental changes. For example at El Kef, Tunisia

18 Ž .d O values suggest a short duration F20 ka

warming above the Ir anomaly, although contempo-raneous changes in the dinoflagellate cyst popula-tions suggest AcoolB taxa dominated this intervalŽ .Brinkhuis et al., 1998 . However, normal tempera-ture preferences may have been of little consequencefor plankton in the curious low-productivity oceansof the earliest Tertiary. The intensity of planktonicextinctions appears to have declined at higher lati-tudes, where the event may also have been a more

Ž .protracted one Pardo et al., 1999 . In the highsouthern palaeolatitudes of New Zealand only 15%of radiolarian species fail to cross the K–T boundaryŽ .Hollis, 1996 .

Clearly, care has to be taken not to confusecorrelation with causation, but the evidence for sud-den extinction and dramatic climatic changes imme-diately above the Ir anomaly in sections like El Kefand Caravaca is compelling evidence for a bolideimpact-triggered mass extinction, with the most dev-astating consequences occurring in equatorial lati-tudes. The effect of the contemporaneous DeccanTraps eruptions is far more difficult to judge. Themajor mid-Maastrichtian events predate the eruptionswhile the later Maastrichtian extinctions are mostobviously related to the rapid eustatic oscillationsimmediately preceding the K–T boundary. Only theglobal warming in the last 0.5 Ma of the Cretaceousmay be ascribed to the Deccan eruptions and thisclimatic event does not coincide with any extinctionevent.

12. Brito–Arctic flood basaltsrrrrrNorth Atlantic Ig-neous Province

Early Tertiary rifting between northern Europeand Greenland produced the youngest of the largeCFBPs, known as the Brito–Arctic Province and the

Ž .North Atlantic Igneous Province Fig. 7 . Volumeestimates range from 2=106 km3 to as high as6=106 km3 if offshore, submarine flows are in-

Ž .cluded in the total Saunders et al., 1997 . Like theSiberian Traps, and perhaps the CAMP, the Brito–Arctic Province consists of several sub-provinceswith distinct eruption histories. The earliest, verybrief eruption interval began around 61.0 Ma in Westand Southeast Greenland and in northwest BritainŽ .Saunders et al., 1997; Storey et al., 1998 . A sec-


ond, voluminous phase of volcanism began around56 Ma and saw the emplacement of basalt flows inEast Greenland, the Faroes and over extensive areasof the North Atlantic margin where they now formseaward-dipping reflectors in offshore seismic linesŽ .Saunders et al., 1997 .

The Faeroe–Greenland volcanics were eruptedwith unusual violence by the standards of floodbasalt provinces, and basaltic tuffs are common in

Ž .the successions Knox and Mortan, 1988 . Pyroclas-tic activity peaked with the development of theBalder Formation in the latest Palaeocene, a majorairfall ash deposit that covers large areas of thenorthern North Sea and west of the Shetland IslandsŽ .Eldhom and Thomas, 1993 . The N. Atlantic wasnot the only site of major pyroclastic activity in the

Ž .later Palaeocene, Bralower et al. 1997 has recordedŽ 5 .a brief phase -10 years of intense explosive

activity in the Caribbean region around 55.0 Ma.The Brito–Arctic volcanicity may coincide with

Žsome dramatic environmental changes but not with.an associated mass extinction . The late Palaeocene

was marked by a gradual shift to the warmer condi-tions that typify much of Eocene time, but the trendwas punctuated by an extraordinarily intense andshort-lived climatic event known as the Late

ŽPalaeocene Thermal Maximum or LPTM Kennettand Stott, 1991; Koch et al., 1992; Robert and

.Kennett, 1994; Norris and Rohl, 1999 . Around 54.9¨Ma both d

18 O and d13C records reveal a brief nega-

tive spike, with the latter showing the developmentof a y3‰ inflexion in only a few thousand yearsfollowed by a AgradualB return to pre-excursion val-

Ž .ues in 120 ka Norris and Rohl, 1999 . This event¨coincides with a temporary increase of deep sea andhigh latitude water temperatures by as much as 78CŽ .Kennett and Stott, 1991 , and an influx ofkaolinite-rich clays into the oceans that is probablythe result of increased weathering in warmer and

Ž .more humid climates Robert and Kennett, 1994 . Amajor immigration event of mammals into NorthAmerica at this time, may also be a response to the

Ž .warming event Koch et al., 1992 . However, theonly significant extinctions were amongst deep-seabenthic foraminifera. These losses appear related tothe development of oxygen-poor bottom watersŽKennett and Stott, 1991; Eldhom and Thomas, 1993;

.Kaiho, 1994 .

Comparison of the timing of the LPTM and vol-canic events has suggested that the two phenomena

Žmay be related Kennett and Stott, 1991; Dickens et.al., 1995 . The initial phase of volcanism was long

before the late Palaeocene event whereas the secondphase appears to have begun around 1 Ma before theevent. The magnitude of the d

13C excursion is toogreat to have caused by volcanic CO , instead2

methane hydrate is widely regarded to have suppliedŽthe light C Dickens et al., 1995; Norris and Rohl,¨

1999; Bains et al., 1999; Katz et al., 1999; Dickens,.1999 , although the eruption of volcanic CO is a2

potential cause of the warming event required totrigger hydrate dissociation. Potentially the LPTMmay also have been terminated by volcanic activity,due to cooling by sulphate aerosols during thewidespread interval of pyrolcastic volcanism in the

ŽNorth Sea region and the Caribbean Beerling and.Jolley, 1998 .

13. Ethiopian and Columbia River flood basalts

The two youngest CFBPs are also the smallest,and they appear to have had little if any impact onthe world’s biota. The Ethiopian Traps consist of0.75=106 km3 of lava which erupted from 31–28

Ž .Ma Courtillot et al., 1999 . Rampino and StothersŽ .1988 linked these eruptions with a protracted lateEocene extinction, but this 34 Ma event considerablypredates the Ethiopian eruptions. The even smallerColumbia River Traps contain 0.17=106 km3 oflava and they have been linked with a mid-Miocene

Žextinction event with which they coincide Coffin.and Eldholm, 1994 . However, this minor event, first

identified in the compilation of Raup and SepkoskiŽ .1984 , is only notable for elevated extinction ratesamongst the ever-vulnerable, deep-sea benthic

Ž .foraminifera which Kaiho 1994 attributes to pro-longed cooling of deep ocean waters. North Ameri-can mammals also suffered extinctions in the late

Ž .Miocene Hallam and Wignall, 1997 , but theselosses were several million years after the peak ofColumbia River eruptions.

14. Discussion

This review has focussed on two key aspects ofthe link between flood basalt volcanism and mass


Ž .extinctions, the correlative timing or lack of it andthe nature of the associated environmental changes.Both aspects suggest an interesting but poorly re-solved relationship.

14.1. A partial correlation?

By far, the most compelling evidence for a linkbetween volcanism and extinctions comes from thecomparison of the ages flood basalt provinces and

Ž .mass extinction events Table 1 . Thus, of 11 majorepisodes of flood basalt formation, five coincideclosely with mass extinctions, and two more, theCaribbean–Madagascar and Brito–Arctic Provinces,coincide with minor extinction events. CourtillotŽ .1994, 1999 has suggested an even better 7-out-of-11correlation, but his compilation includes the spuriouslink between Parana–Etendeka volcanism and an´

Ž .end-Jurassic mass extinction see above , togetherwith a so-called end-Palaeocene Amass extinctionB.The correlation with episodes of oceanic anoxia is

Ž .even better 7 out of 11 , although only for four ofthese is the anoxia regarded as the proximate cause

Žof marine extinctions the end Permian, early Toar-.cian, end Cenomanian and late Palaeocene events .

The volcanism–extinction link is best seen for thefour mid-Phanerozoic events that began with theeruption of the Emeishan flood basalts and endedwith the formation of the Karoo–Ferrar Provinces.However, for the first three of these examples, theeruptions appear to have begun a short but signifi-

cant interval after the extinctions. The youngerCaribbean–Madagascar volcanism may have simi-larly post-dated the end-Cenomanian extinction.Clearly there is still a need for more precise datingof these provinces, but the slightly non-synchronouscorrelation suggests that the onset of eruptions orperhaps the interval immediately preceeding this isassociated with the most damaging environmentalchanges. One such change may be widespread upliftprior to eruption. This has been suggested for theCAMP, with the consequent reduction in shallowmarine habitat area being a potential cause of extinc-

Ž . Ž .tion Hallam, 1990 , although Hames et al. 2000note there is no evidence for pre-eruption doming inthis case. Major sea-level fall may also have been afactor in the end-Guadalupian and end-Triassic massextinctions, but not for the end-Permian and early

Ž .Toarcian events Hallam and Wignall, 1999 . How-ever, it is difficult to envisage how these globalsea-level changes can be attributed to purely regionaldoming.

Post-Jurassic links between volcanism and extinc-tions are substantially more tenuous. Only the Dec-can Traps coincide with a mass extinction, but thecelebrated Chicxulub impact was probably the primeculprit in this case. Nonetheless, three of the largeigneous provinces formed in the Cretaceous–

ŽPalaeocene interval the Ontong–Java, Caribbean–.Colombian and Brito–Arctic Provinces correlate

with remarkably similar oceanicrclimatic events,namely the development of widespread marine anoxia

Table 1Comparison of environmental and biotic events that were contemporaneous with the eruption of large igneous provinces

Flood basalt provinces Coeval mass extinction Global C isotope excursions MarineŽ .warming ‰ anoxia

Emeishan–Panjal Volcanics end-GuadalupianSiberian Traps end-Permian 6 y6.0 6

CAMP end-Triassic 6 6?Karoo–Ferrar Traps early Toarcian 6 y3.0 and then q4.5 6

Parana–Etendeka TrapsÓntong–Java Plateau 6

Caribbean–Colombian volcanism end-Cenomanian 6 q2.0 6

Deccan Traps end-Cretaceous 6 y2.0Brito–Arctic Flood Basalts foraminifera extinctions only 6 y3.0 6

Ethiopian TrapsColumbia River Flood Basalts


and equally dramatic global temperature increases.Perhaps the most intriguing aspect of these events istheir relatively small impact on the biosphere.

14.2. Volume and Õiolence of eruptions

Estimating the original volumes of lava in floodbasalt provinces is rendered difficult due to subse-quent erosion, partial destruction during continentalcollisions or burial beneath passive margin sedimen-tary wedges. Nonetheless, comparison of best-esti-mate volumes with extinction intensity suggests that

Ž .there is no correlation Fig. 9 . Clearly, the lavavolumes alone are not a key factor in mass extinc-tions, but the rapidity of eruption may be. Unfortu-nately, there has been little documentation of thesizes of individual flows within flood basalt

Ž .provinces, Courtillot’s 1990 widely quoted 10 000km3 figure is only a guess.

Several recent studies have highlighted that theproportion of pyroclastic flows in flood basalt

provinces has probably been significantly underesti-mated. Thus, it could be argued that the high propor-tion of pyroclastics within the Siberian Traps indi-cates that violence of eruption is a key factor inwreaking environmental damage. However, a counterexample is readily found in the Brito–ArcticProvince, which also contains a high proportion ofpyroclastics, and yet it does not coincide with a massextinction. If the formation of LIPs is associated withfrequent explosive eruptions capable of injectinggases into the stratosphere then evidence should befound for the resultant effect: sulphate aerosol-in-duced cooling.

Many previous studies have emphasised the globalcooling effects of sulphate aerosols as the prime

Žcause of volcanogenic extinctions Rampino and.Stothers, 1988; Courtillot, 1994, 1999 . This sugges-

tion accords with observations following historicalŽ .eruptions Devine et al. 1984 , but the evidence for

cooling from the geological record is, at best, tenu-Ž .ous. Kozur 1998 provides some of the little evi-

Ž .Fig. 9. Comparison of generic extinction percentages, from Sepkoski 1996 , with estimated original volumes of coeval igneous provinces,showing no correspondence between the two.


dence available for the end-Permian mass extinctionŽ .but see the alternative explanations discussed above .The failure to find evidence for cooling events trig-gered by volcanism is probably because the short

Žresidence time of aerosols in the atmosphere -5. Žyears is insufficient to cause climatic change or

.extinctions .

14.3. Volcanic CO emissions2

In contrast to the dearth of evidence for cooling,rapid global warming coincides with at least 6 of the

Ž .11 intervals of volcanic trap formation Table 1 .The eruption of large volumes of CO is widely2

Žregard as the most likely cause. In three cases end.Permian, Toarcian and late Palaeocene , the warming

events coincide with rapid negative d13C shifts of

amplitudes that are too great to be entirely at-tributable to the input of volcanic CO into the2

atmosphere. For these cases the volcanism is inter-preted to have initiated warming that triggered the

Ž .dissociation of isotopically light methane hydrateswith a resultant exacerbation of the warming trendŽe.g. Erwin 1993; Dickens et al., 1995; Bowring et

. Ž .al., 1998; Hesselbo et al., 2000 . Kerr 1998 hassuggested a similar cause-and-effect for volcanism atthe Cenomanian–Turonian boundary although nei-ther this interval, nor the other intervals of floodbasalt province formation, are associated with thesame negative spike in d

13C values.Ž .Courtillot 1999, p. 98 has suggested that Athe

w xmagnitude of the biological effects of volcanismwill depend on a large number of factors: the ar-rangements of the continents, the sea level and cli-mate at the time of the eruptions, the total amplitudew xviolence? , duration and number of individual events,the closeness in time of these events, and so on.BDisentangling all these factors will clearly be amajor challenge.

14.4. A lack of pre-Permian LIPs?

CFBPs have formed on average once every 30 Masince the Middle Permian, and several late Protero-zoic examples are known, now eroded down to their

Žintrusive roots Pelechaty, 1997; Ernst and Buchan,.1997; Li et al., 1999; Barovich and Foden, 2000 ,

but no examples are known from the 240 Ma spanfrom the Cambrian to the Middle Permian. If LIPs

are the product of plume eruptions, then their ab-sence from most of the Palaeozoic implies funda-mentally different mantle dynamics in this interval.Alternatively, and more plausibly, the absence ofPalaeozoic CFBPs may reflect the lack of majorrifting episodes in this interval, an inference which

Žsuggests that flood basalts are the product not the.cause of rifting as postulated in the models of White

Ž . Ž .and McKenzie 1989 and Anderson et al. 1992 .Of the several major mass extinction events that

occur prior to the end-Guadalupian crisis, only theŽ .late Devonian Frasnian–Famennian boundary event

has been linked with a possible volcanogenic cause.The crisis is marked by the loss of diverse marineinvertebrate groups, especially in low palaeolati-tudes, and the subsequent spread of cool water taxa

Žin the immediate post-extinction interval McGhee,.1996; Copper, 1998 . The spread of anoxic waters

and rapid sea-level oscillations are further features ofŽ .this interval Becker and House, 1994 . Volcanism

has been implicated as a contributory cause of theseenvironmental changes, with eruptions postulated to

Žoccur at either an intraoceanic site Becker and.House, 1994 , or in a major continental rift preserved

Ž .in the Ukraine Racki, 1998 . The volcanism associ-ated with this Pripyat–Dnieper–Donet rift is volu-

Ž 3.metrically minor -10 000 km , compared withtrue LIPs, but its environmental effect may havebeen proportionally greater due to the dominance of

Ž .pyroclastics Wilson and Lyashkevich, 1996 . Betterdating of the Ukraine volcanism is needed in order toascertain its potential relationship with theFrasnian–Famennian crisis.

15. Conclusion

Ž .As Courtillot 1999 remarked, there is a Are-markable correlationB between the age of flood basaltprovinces and mass extinction. However, his conclu-

Žsion that Athe correlation is almost perfectB Courtil-.lot, 1999, p. 97 is too optimistic, particularly if

Palaeozoic mass extinctions are considered. This re-view reveals the following relationships:

Ž .1 Of the 15 major Phanerozoic extinctions, onlysix coincide with major episodes of volcanicity.However, all extinction events of the last 300 Macoincide with LIPs, with. the best correlation occur-ring for the 4 mid-Phanerozoic event that began with


the end-Guadalupian mass extinction. Subsequent ex-tinctions have been less intense, with the exceptionof the K–T event whose magnitude is likely due tothe effects of bolide impact not volcanism.

Ž .2 For many LIPs, the onset of eruptions, or theinterval immediately before this, appears to coincidewith the extinction interval, perhaps suggesting thatvolcanicity triggered harmful environmental change.

Ž .3 The eruption volume of flood basalt provincesŽ .is unrelated to extinction intensity Fig. 9 , and

neither is their duration, which once again argues fortheir role in triggering further change, such as gashydrate release.

Ž .4 Episodes of global warmingrmarine anoxiashow the best correlation with flood basalt volcanismŽ .Table 1 . Thus, all post-Palaeozoic oceanic anoxic

Ževents coincide with intervals of LIP formation e.g.Keith, 1982; Sinton and Duncan, 1997; Kerr, 1998;

.Jenkyns, 1999 . This connection may be connectedwith volcanic CO emission, particularly if the2

warming effect is exacerbated by gas hydrate disso-ciation. However, it is an unresolved paradox thatthese warmingranoxia events are associated with

Žboth major mass extinctions e.g. the end-Permian. Ževent , modest extinctions, e.g. the end-Cenomanian. Ževent , very minor extinctions e.g. the late

. ŽPalaeocene and no extinctions at all e.g. the Early.Cretaceous Selli Event .

Ž . Ž .5 The duration of volcanogenic? oceanicanoxia varies greatly. Thus, the P–Tr marine anoxicevent persisted for several million years, whereas theLate Palaeocene thermal maximum lasted for only100 ka. The efficiency of C sinks, and their ability todrawdown atmospheric CO levels may control the2

duration of Arunaway greenhouseB intervals and thustheir impact on the world’s biota. This efficiencyappears to have undergone a secular increase in thepast 300 Ma, with early Mesozoic volcanism havingsubstantially more environmental impact than Ter-tiary volcanism.

Acknowledgements

I thank Mike Widdowson and Tony Hallam forcomments on an earlier version of this manuscript,and Andy Saunders and Olav Eldholm for valuable,constructive criticism.

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Paul Wignall is a Reader in Palaeoenvi-ronments at the University of Leedswhere he has been a resident since 1989.Prior to this, his PhD studies at theUniversity of Birmingham, and subse-quent post-doctoral studies at the Uni-versity of Leicester, were focussed onthe palaeoecology of black shales. Hiscurrent research interests include thestudy of marine environmental changeduring mass extinction intervals, mostnotably during the end-Permian and

end-Triassic crises. In recent years, he has been documenting thenature and extent of oxygen-deficiency during the Permian–Tri-assic boundary interval in various far-flung corners of the globe.This research has led him to take an intense interest in thepotential environmental impact of large igneous provinces andcontinental flood basalts.

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Large igneous provinces and mass extinctions