OPEN Thermal erosion of cratonic lithosphere as a ... · Thermal erosion of cratonic lithosphere as a potential trigger for mass-extinction ... Cratonic lithosphere as a potential

1SCIENTIFIC REPORTS | 6:23168 | DOI: 10.1038/srep23168

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Thermal erosion of cratonic lithosphere as a potential trigger for mass-extinction

The temporal coincidence between large igneous provinces (LIPs) and mass extinctions has led many to pose a causal relationship between the two. However, there is still no consensus on a mechanistic model that explains how magmatism leads to the turnover of terrestrial and marine plants, invertebrates and vertebrates. Here we present a synthesis of ammonite biostratigraphy, isotopic data and high precision U-Pb zircon dates from the Triassic-Jurassic (T-J) and Pliensbachian-Toarcian (Pl-To) boundaries demonstrating that these biotic crises are both associated with rapid change from an initial cool period to greenhouse conditions. We explain these transitions as a result of changing gas species emitted during the progressive thermal erosion of cratonic lithosphere by plume activity or internal heating of the lithosphere. Our petrological model for LIP magmatism argues that initial gas emission

2 became the

not associated with climatic and biotic crises comparable to LIPs emitted through cratonic lithosphere.

2 concentrations1,2

2 from volcanic basalt prov-inces3 and/or from the degassing of carbonaceous or organic-rich sediments by sill and dyke intrusions4,5

second scenario invokes a short period of global icehouse conditions caused by degassing of large volumes of vol-canic SO2, atmospheric poisoning, cooling, and eustatic regression6,7. Although both hypotheses are compatible with massive volcanic degassing, they must also be able to explain the paleontological record in marine and conti-nental stratigraphic sections. Mass extinction events are recorded in various marine and continental sedimentary sections distributed around the world. One major challenge is to correlate these sections in order to obtain a global “picture” required to understand global climate change associated to mass extinction. Marine versus conti-nental sedimentary sections do not necessarily record the same processes neither the same time interval. In par-ticular, LIPs volcanic activity related to major periods of extinction (Siberian Trap, CAMP, Karoo-Ferrar, Deccan) is restricted to continental settings while key observations to demonstrate global mass extinction is recorded in marine environments. It requires, therefore, precise and accurate time constraints to link volcanic activity with the change in the paleontological record, precision at the 100 ka level, which is only achievable using high preci-sion U/Pb geochronology on zircon8,9. Here, we review data from the stratigraphic record of the Triassic-Jurassic (T-J) and Pliensbachian-Toarcian boundaries combined with geochronological data8–10 in order to establish the

alternatives to explain the climatic and chemical changes associated to these sequences of events are discussed.

Sequence of events associated to the Triassic-Jurassic boundaryDetailed ammonite biostratigraphy7,11 combined with the carbon and oxygen stable isotope records1,12,13 and zircon geochronology8,9 allows reconstruction of the sequence of events associated with biological crisis observed

1Institute of Earth Sciences, University of Lausanne, Géopolis, 1015 Lausanne, Switzerland. 2Muséum National

Paris, France. 3Institute of Earth Surface Dynamics, University of Lausanne, Géopolis, 1015 Lausanne, Switzerland. 4Department of Geosciences, Princeton University, 219 Guyot Hall, Princeton, New Jersey 08544, USA. 5Earth & Environmental Sciences, University of Geneva, Rue des Maraîchers 13, 1205 Geneva, Switzerland. Correspondence and requests for materials should be addressed to S.P. (email: [email protected])

Received: 14 October 2015

Accepted: 29 February 2016

Published: 24.03.2016

OPEN

Jean Guex1, Sebastien Pilet1, Othmar Müntener1, Annachiara Bartolini2, Jorge Spangenberg3, Blair Schoene4, Bryan Sell5 & Urs Schaltegger5

mailto:[email protected]

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et al.14, between end-Triassic extinction (ETE) marked by the major extinction of the Triassic ammonoids (indicated by the last occurrence (LO) of Choristoceras crickmayi) and the earliest CAMP volcanism7–9 -ized by a strong negative excursion of δ 13Corg recorded in New York Canyon (Nevada, USA) and worldwide7,8 correlated to a initial negative excursion of δ 13Cwood measured in fossil leaf and wood from East Greenland1,12,15

-sion which predates or is concomitant with environmental changes at the Rhaetian – Hettangian transition is supported by various stratigraphic evidence: (1) At New York Canyon, the end Triassic sediments are character-ized by a bed packed with small shallow water bivalves indicating low sea level conditions; (2) In northern Peru (Chilingote, Utcubamba Valley) an erosional tempestite horizon located just below Odoghertyceras ammonite (early Jurassic) bearing sediments indicates this section was above storm wave base in the topmost Triassic16; (3) A pronounced increase in the proportions of terrestrial versus marine palynomorphs during the end-Triassic interval is observed in various sedimentary sequences from the Alps17–19; (4) In England, wave-ripple structures and mudcracks horizons, evidencing shallowest water depth conditions and sporadic emersions, are observed just few centimeters below the end-Triassic negative δ 13C excursion20

δ 18O values as measured in Oysters from Lavernock (UK)13,19.Psiloceras

spelae90 to 200 kyr, based on sedimentary rates constrained by U-Pb zircon geochronology from ash beds in Northern Peru8,11 δ 13Corg, decrease of δ 18

warming potentially associated with large volcanic CO21,2,12,15

warming period is associated with an abrupt change in plant diversity observed in Greenland1,15 and with a sec-ond negative δ 13C recorded in the Hettangian Psiloceras planorbis beds (coeval with ) that postdate the ETE (see supplementary information).

Sequence of events associated to Pliensbachian-Toarcian boundaryEnvironmental perturbations related to the early Jurassic Pliensbachian-Toarcian boundary have been associated for some time with the onset of the Karoo-Ferrar large igneous province21

precision U-Pb dating on zircons10,22 which indicate that major sills intruded into organic-rich sediments in the Karoo basin are correlated with the Toarcian Oceanic Anoxic Event (OAE) recorded in the marine sedimentary

extinction, preceding the Toarcian OAE23, is marked by an important diversity drop (disappearance of 90% of the ammonite taxa) associated with a generalized sedimentary gap linked to a marked regression event in

24

the growth of ice caps25–28, and was marked by important emersion topography observed on seismic images of the North Sea24, evidence of polar ice storage29

Nevada (USA)30 and Ururoa-Kawhia area, New Zealand31) deposited in open marine sediments containing Late Pliensbachian ammonites (Nevada) and below sediments containing Early Toarcian ammonites (Nevada and

δ 18O data on belemnites27.

Figure 1. Stratigraphic and isotopic correlations for the Triassic-Jurassic boundary. (a) U-Pb ages on zircon from ash beds embedded in marine stratigraphic sections from N. Peru and Nevada8. (b) Standard ammonite zones for the late Triassic- early Jurassic. (c) Duration of CAMP volcanism in North America and Morocco (Argana Basin)9. (d) Ammonite biodiversity around the T-J boundary. (e) δ 13Corg from New York Canyon (Nevada, USA)11. (f) δ 13Corg recorded in stratigraphic sequences from East Greenland and pCO2 estimated for the late Triassic-early Jurassic1,2,15. (g) δ 18O measured on Oysters from Lavernock (UK)13. (h) Sea level variations8. (i) Variations of the temperature15. (j) Proposed model for the decoupling between sulfur emissions related to the thermal-erosion of the continental lithosphere following by CO2 degassing associated to CAMP volcanism. Notes: ETE: End Triassic Extinction. MPT: Early Jurassic major plant turnover1. LO: Last Occurrence. FO: First Occurrence. Neo & Ppl: Neophylites & Psiloceras planorbis ammonite species respectively. (See supplementary information for additional details about the correlation between marine and continental sections).


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of black shales associated with the Toarcian OAE32 -oeustatic sea-level rise associated to rapid change from cold to warm climatic conditions33

Toarcian OAE in S. Peru has been dated at 183.22 ± 0.25 Ma, coinciding with the oldest Karoo sill and lava cur-rently dated (granophyre sill intruded in the Tarkastad Formation date of 183.014 ± 0.054 Ma10 and granophyre

± 0.045 Ma2234–36 but the ammonites were only slightly

or are concomitant with evidence of environmental change at the Rhaetian – Hettangian and the Pliensbachian–

duration of these events (< 1 -cales associated with heat-induced changes in lithospheric buoyancy37. Moreover, such long-term processes seem

with large initial CO2 degassing associated with LIPs as proposed to explain the T-J and Pl-To mass extinctions (e.g. refs 1,2) or by volatile-release (CO2, CH4, Cl2) from deep sedimentary reservoirs during contact metamor-phism associated to dykes and sills intrusion4,5 because massive CO2 degassing is expected to produce super

-6,38. Measurements of

sulfur content in melt inclusions or in clinopyroxene demonstrate that S contents in LIP basaltic melts could be high (> 1000 of recycled oceanic or continental crustal lithology in the mantle source of LIPs or crustal contamination3,5,39,45 has been proposed to explain these high S contents regarding that the melting of the depleted MORB mantle (119 ppm of S46) alone seems unable to explain such high values. Here we evaluate the hypothesis that initial ther-mal erosion of the cratonic lithosphere due to emplacement of the CAMP and Karoo-Ferrar volcanic provinces led to initial pulses of sulfur causing global cooling and eustatic regression, which was followed by warming/transgression associated with the progressive increase of CO2 in the atmosphere associated to LIPs emission and metamorphic reactions in sedimentary basins4,5.

Cratonic lithosphere as a potential sulfur reservoirPetrological constraints on primary magmas indicate that the mantle is hotter and melts more extensively to

47

garnet-bearing peridotitic sources at high pressures (5–6 GPa) and anomalously high mantle potential tem-peratures (Tp) of > 1,600 the Karoo-Ferrar province48 while lower Tp (1450 °C ± 50 °C) and lower pressures of melting (> 2 GPa) were estimated for tholeiitic lavas from CAMP and Ferrar large igneous province49. Available data suggest that the Karoo-Ferrar and CAMP have been emitted on top of thick lithosphere50. First, the eastern and the southern extents of the CAMP were located on top of Archean and Proterozoic cratons, while the Karoo LIP was erupted on and around the Kaapvaal craton. Shear wave velocities (Vs), suggest that the Karoo and CAMP areas were

Figure 2. Stratigraphic and isotopic correlations for the Pliensbachian-Toarcian boundary. (a) Ammonite zones for the late Pliensbachian – early Toarcian. (b) U-Pb ages from ash beds collected at Palquilla (Peru)10. (c) U-Pb ages of the Karoo and Ferrar sills and lavas10,22. (d) Fluvial conglomerate deposits in Nevada and New Zealand30,31. (e) Variations of the biodiversity of ammonites during the Pliensbachian- Lower Toarcian23. (f) Variations of δ 13C in the Peniche section80. (g) Variations of Hg/TOC (TOC: total organic carbon) in the Penische section72. (h) Variations in δ 18O of belemnites (color-coded as function of their Ammonite zone location) and relative changes in temperature27. (i) Temperature variations around the Pliensbachian-Toarcian boundary. (j) Sea level change interpreted based on δ 18O variation and on stratigraphic constraints. (k) Proposed model for the decoupling between sulfur emissions related to the thermal-erosion of the cratonic lithosphere following by CO2 degassing associated to Karoo-Ferrar dike swarm emplacement and volcanism. Ammonite zones: THOU: ; VAR: Varabilis; BIF: Bifrons; FAL C: Falciferum; TEN: Tenuicostatum; Spin = Spinatum; Marg = Margaritatus. Notes: Toarcian OAE: Toarcian Oceanic Anoxic Event.


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underlain by thick lithosphere (> 200 km) prior to continental break up51–53, which postdates the LIP emplace-ment54

55. A thick lithosphere exceeding 200 km beneath the Kaapvaal craton is also documented by kimberlite-borne mantle xenoliths56

thickness of the lithosphere beneath the Ferrar LIP is probably similar, given that the Ferrar LIP emission was linked to the Wilkes Land craton57. Taking into account estimates of the thickness of continental lithosphere beneath Karoo-Ferrar and the CAMP area, the melting conditions estimated for LIP’s meimechites and tholeiitic lavas48,49 point out the fundamental role of the lithospheric mantle to produce these lavas. Various geochemical and isotopic studies on CAMP58–60 and Karoo-Ferrar magmas61–65 indicate that continental lithospheric mantle

explain the melting of the lithospheric mantle65: the arrival of a thermal plume initiating heating and subsequent partial melting of the subcontinental lithosphere3, or internal heat production of mantle insulated by continen-tal lithosphere allowing for partial melting of the upper asthenospheric and continental lithospheric mantle to produce LIPs magmas49,63,66–68

critical to understand the volatile budget associated with LIPs while studies of the composition of the Kaapvaal 69

base of the subcontinental lithospheric mantle from the Archean to the Proterozoic69.

internal heating of the subcontinental mantle could release sulfur from the cratonic roots. (I) Low degree melts

70

lithologies will be remobilized by the progressive percolation of low degree melts from the subcontinental lith-

directly reach the surface, but enrich progressively shallower levels of the lithospheric mantle, levels which could

2

lavas and carbonatites associated with translithospheric fracturing as observed in various LIP areas71

-tion of large volumes of magma into sedimentary basins releases additional sediment-derived greenhouse gases (CO2, CH4, …) produced during contact metamorphism4,5

emphasizing an important role of the lithospheric mantle in the petrogenesis of Karoo-Ferrar and CAMP LIPs,

precursor phase of magmatism associated to the Karoo-Ferrar LIP is supported by recent mercury data deter-mined in marine sediments72

Figure 3. Pressure-Temperature diagrams for the condition of melting of dry-peridotite compared to the thickness of the lithosphere before thermal erosion (> 200− 230 km54,69

CAMP and Karoo-Ferrar CFB are estimated by Herzberg and Gazel4748 and CAMP

tholeiites49 81 and 70 GPa, the yellow dashed

line indicates extrapolation to higher P-T conditions (See supplementary information for details about the melting condition in an upwelling plume).


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ratios correlated with the negative carbon-isotope anomalies associated with the Pliensbachian-Toarcian bound-ary and Toarcian OAE observed in Peniche, Arroyo Lapa, Mochras and Bornholm sections are indicators for two distinct episodes of Hg release to the atmosphere by volcanic activity72 in line with our sequence of two distinct magmatic events.

To estimate the potential amount of S released to the atmosphere during the initial phase of lithosphere ther-mal erosion/heating, we have calculated the amount of sulfur present in the lower 25 km of the lithosphere and assuming a surface equivalent to half the area covered by Karoo-Ferrar or CAMP LIPs. Based on garnet chemistry

-tions of melt metasomatized peridotites56. Assuming that the basal part of the lithosphere was composed of 40% depleted peridotite (119 ppm S), 55% metasomatized peridotite (300 ppm S) and 5% pyroxenite (800 ppm S), we obtain a potential sulfur content that could be released to the atmosphere (calculated in equivalent SO2) of ~45,000 Gt and ~210,000 Gt of SO2 for the case of the Karoo-Ferrar and CAMP, respectively (see supplementary information for details). Our estimates for potential sulfur release from the lithospheric mantle are 5 orders of magnitude higher than the mass estimated for Laki (~122 MT SO2

73) or Pinatubo (~20 Mt SO274) eruptions.

3 has shown that gas release from plume mate-

Figure 4. Schematic sections illustrating the thermal heating of the lithospheric mantle associated to the arrival of a thermal plume beneath a thick cratonic lithosphere (panels (a–c)) or linked to the increase of mantle temperature through supercontinent thermal insulation (panels (d,ecomposed of melt-metasomatized peridotite69. Panels (a,d) illustrate the initial step of melting, panels (b,e), the precursor step of magmatism suggested by Ernst and Bell71 while panels (c,f) show schematic sections of

suggested by Heinonen et al.65 for the magma generation of Karoo-Ferrar CFB.


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gas released from the cratonic lithosphere. We suggest here that the initial gas release to the atmosphere is not solely composed by gas from the asthenosphere itself (CO2/HCl3), but additional sulfur gases produced from

period. Studies of recent catastrophic eruptions demonstrate that large SO2

2 release into the atmosphere demonstrate that it is not the magnitude of SO2 emission, which is critical to produce cooling, but the frequency and the duration of SO2 emission38 -dictions about the heights reached by the volcanic clouds, we provide feasibility tests. Assuming that only 20% of the sulfur from the metasomatized lithospheric mantle is released to the atmosphere during a period of 100,000

Mt/yr for the case of Karroo and CAMP, respectively. Timmreck et al.77 have modeled that sulfur emissions of ~100 times the mass of Pinatubo (corresponding to our estimate of S emitted for a period of 5 to 20 years) creates an average temperature decrease of − 3.5 °K lasting for 9–10 years. Robock75 hypothesized that a series of volcanic eruptions could initiate a longer-term cooling period. Schmidt et al.38 support this hypothesis but indicate that SO2 release could cause biotic crisis only if such gas release is high and sustains for several centuries. Multiple SO2 pulses lasting for several thousands of years as suggested here

For example, if the Pliensbachian–Toarcian boundary is associated to ammonite and planktonic massif turn over,

et al.382 degassing could be variable depending on the eco-

system and location. Since the residence time of sulfur in the atmosphere is limited6,73,75,77

of CO2, CH4, Cl2-

ontological records. We hypothesize that this last stage is responsible for the second extinction observed at the T-J

Proterozoic continental lithosphere is therefore a major component in the timing of mass extinction processes

continental LIPs erupted on cratonic lithosphere such as the Siberian Trap, CAMP, Karoo-Ferrar and Deccan. -

tant control on the consequences on life at the Earth’s surface and is a viable hypothesis in addition to external causes such as asteroid impacts as proposed to explain the Cretaceous-Paleogene crisis78). Indeed, higher sulfur contents in CAMP and Deccan magmas, with respect to the Paraná LIP have been reported39

magmatic sulfur contents were explained by the presence of slab- metasomatized mantle in the source of CAMP and Deccan magmas39. While the CAMP and Deccan LIPs were associated to major extinctions, the emission of the Paraná - Etendeka province was not79

Figure 5. Illustration of mass extinction intensity associated with continental LIPs and oceanic plateaus against geological time. estimated volume (in km3) of emitted lavas82–84

the underlying lithosphere (lithosphere ages from ref. 50).


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such as the Ontong Java or the Caribbean plateaus, did not cause climatic and biotic crises as large as those of the Karoo-Ferrar or the CAMP despite having erupted similar volumes of basaltic rocks.

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Author Contributions

involved in the preparation of Figures 1 and 2 summarizing the sequences of events associated with the T-J and Pl-To boundaries. S.P and O.M. developed the petrological model for thermal erosion of the lithosphere and

J.G., U.S. and Bl.S. wrote the text. All authors reviewed and approved the manuscript.

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Supplementary Information: Guex et al. Sci. Rep. 6, 23168 (2016) __________________________________________________________________________________________________

1

SCIENTIFIC REPORTS !

Thermal erosion of cratonic lithosphere as a potential trigger for mass-extinction

!

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1*,%Othmar%Muntener

1,%Annachiara%Bartolini

2,%Jorge%Spangenberg

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5%Section%of%Earth%&%Environmental%Sciences,%University%of%Geneva,%Rue%des%Maraîchers%13,%1205%Geneva,%Switzerland%

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SUPPLEMENTARY INFORMATION

PART I. Supplementary discussion concerning the stratigraphic correlation associated with the timing of sea level changes, δ13C, δ18O, pCO2 variations, paleotemperatures, and the age of the onset of the CAMP-related basaltic

extrusions One$ of$ the$ major$ challenges$ for$ the$ worldwide$ stratigraphic$ correlation$ of$ the$ Triassic$ Jurassic$ (T9J)$ and$Pliensbachian$ Toarcian$ (Pl9To)$ boundaries$ is$ to$ correlate$ the$ continental$ and$marine$ sedimentary$ record.$We$present$a$synthetic$stratigraphic$correlation$for$marine$and$continental$settings$with$the$T9J$correlation$shown$in$figure$1$of$the$main$article.$$$$Rhaetian - Hettangian boundary !!1)!Correlation!between!marine!and!continental!sections!!!

Figure$ SI91$ illustrates$ the$ correlation$ between$ marine$ and$ continental$ sections$ of$ the$ Rhaetian9Hettangian$boundary.$ The$ synthesis$ of$ the$ Late$ Rhaetian$ and$ Hettangian$ ammonite$ distribution86$ allows$ for$ a$biostratigraphic$ correlation$ of$marine$ environments.$ The$ biostratigraphically$ resolved$ carbon$ isotope$ data87988$can$be$related$to$the$continental$carbon$isotope$record$from$East9Greenland89990$(Fig.$SI91).$U9Pb$data$on$zircon$from$ ash$ beds$ embedded$ in$ the$ Late$ Rhaetian$ 9$ Early$ Hettangian$ of$ marine$ sequences$ in$ Northern$ Peru$ and$central$Nevada$(USA)$provide$absolute$time$constraints$for$the$Rhaetian9Hettangian$transition$that$can$be$linked$to$the$onset$of$CAMP$basalt$activity86,91.$Schoene$et$al.$(ref.$91)$have$demonstrated$the$synchroneity$between$the$eruption$of$ the$oldest$North$Mountain$Basalt$ flows,$and$the$onset$of$ the$ index$ fossil$of$ the$base$of$ the$ Jurassic,!Psiloceras!spelae$and$end9Triassic$biological$crisis$recorded$in$marine$sediments.$$!An$open$question$is$whether$the$end9Triassic$extinction$(ETE)$defined$by$Olsen$et$al.$(ref.$92)$in$the$continental$environment$represents$the$same$event$marked$in$the$oceans$by$the$disappearance$of$Triassic$ammonites.$Olsen$et$al.$(ref.$92)$and$Blackburn$et$al.$(ref.$93)$suggested$that$ETE$is$marked$by$a$dramatic$turnover$in$fossil$pollen,$spores$(sporomorphs),$and$vertebrates$observed$in$early$Mesozoic$basins$of$eastern$North$America.$However,$the$stratigraphy$of$the$Rhaetian$in$the$Newark$supergroup$and$Argana$basin$is$probably$incomplete$and$contains$one$or$several$hiatuses.$Kozur$and$Weems$(refs$94,$95)$were$ first$ to$correlate$ the$conchostracan$biostratigraphy$of$the$ Newark$ Supergroup$with$ the$ Germanic$ Triassic.$ They$ concluded$ that$most$ of$ the$ Rhaetian$ in$ the$ Newark$


2

Basin$ and$ elsewhere$ in$ the$Newark$ Supergroup$ is$missing$ (see$ also$ Lucas$ et$ al.;$ ref.$ 96).$ Gallet$ et$ al.$ (ref.$ 97)$reached$similar$conclusions$on$the$basis$of$paleomagnetic$data.$Therefore,$the$LO$of$Patinasporites!interpreted$as$an$ End$ Rhaetian$ biostratigraphical$ marker93$ could$ potentially$ occur$ in$ the$ lowermost$ Rhaetian98,99.$ Such$ an$interpretation$ is$ supported$ by$ the$ palynological$ study$ of$ Tethyan$ sections$ around$ the$ Norian$ Rhaetian$boundary99.$The$palynological$assemblage$including$Patinasporites$is$typical$of$the$Norian9$Rhaetian$transition99,$and$ Patinasporites$ is$ known$ only$ at$ very$ base$ of$ the$ Rhaetian100.$ This$ suggests$ that$ the$ Newark$ fern$ spike$ is$possibly$ lowermost$ Rhaetian$ in$ age$ rather$ than$ Rhaetian9Hettangian99,100$ (see$ Wotzlaw$ et$ al.,$ ref.$ 101,$ for$ a$discussion$of$the$duration$of$the$Rhaetian).$In$conclusion,$it$is$not$clear$whether$the$end9Triassic$biological$crisis$as$recorded$in$the$Newark$basin$by$Olsen$et$al.$(ref.$92)$and$Blackburn$et$al.$(ref.$93)$represents$the$same$event$recorded$in$other$marine$sequences,$such$as$in$Northern$Peru$and$central$Nevada$(USA),$but$it$does$not$modify$the$conclusion$of$figures$1$and$SI91,$which$shows$that$ETE$is$associated$to$a$clear$first$negative$peak$in$δ13C$and$is$concomitant$or$slightly$predates$the$onset$of$CAMP.$$

$Fig. SI-1. Correlation of Rhaetian‐Hettangian marine sections (Levanto, New York Canyon, Lavernock) with continental records (Newark supergroup, East Greenland). These correlations are constrained by ammonite biostratigraphy, δ13Corg, δ13Cwood from the East Greenland continental record, and U/Pb zircon ages from Levanto and NYC ash beds (red and blue stars; ref. 91) and from zircon from Newark basin lavas (green stars, column 6: refs 91,93 (data from ref. 93 are indicated with an *). The numbers 1 to 3 highlight the proposed correlations between carbon isotope excursions 88. Data references: (1) Fundy Basin – ref. 91; recalculated with new Earthtime spike constants reported in ref 93. (2) Levanto (N. Peru) – refs 86,91. (3) New York canyon – ref. 88. (4) East Greenland – refs 89,90. (5) Lavernock (SW England) – ref. 106. (6) Newark basin – refs 93,102. $2)!Relation!between!extinctions!and!onset!of!the!CAMP!$A$substantial$increase$of$atmospheric$pCO2$interpreted$as$a$consequence$of$important$volcanic$CO2$emissions$has$been$documented$ in$East9Greenland89,90$and$ in$ the$Newark$Basin102$using$stomatal$analysis$of$ fossil$Ginkgoales$leaves$and$pedogenic$carbonates$interbedded$with$volcanics$of$the$(CAMP)$respectively.$$The$analysis$of$Schaller$et$ al.$ (ref.$ 102)$ shows$ that$ the$main$ pCO2$ increase$ appears$ at$ the$ end$ of$ CAMP$ volcanism$ (Fig.$ SI91).$ Precise$correlations$between$crucial$fossil$groups$and$the$δ13Corg$record$for$Rhaetian9$Hettangian$sections$suggests$that$the$Major$Plant$Turnover$defined$by$McElwain$et$al.$(ref.$89)$is$not$correlated$to$the$d13Corg$negative$excursion$observed$ at$ the$ end$ of$ the$ Rhaetian$ (ref.$ 88),$ but$ to$ a$ second$ negative$ δ13C$ peak$ recorded$ in$ the$ Hettangian$Psiloceras!planorbis$beds$(coeval$with$P.pacificum$beds)$occurring$about$600$kyr$after$the$main$ETE.$The$base$of$that$ second$major$ negative$ excursion$ correlates$with$ the$ lower$ part$ of$ the$main$ negative$ carbon$ excursion$ of$Hesselbo$ et$ al.$ (ref.$ 103),$ attributed$ to$ the$ Psiloceras! planorbis$ beds.$ The$ second$ extinction$ event,$ mainly$continental,$ is$ observed$ in$ the$ P.! pacificum$ beds88.$ Note$ that$ the$ palynological$ record$ in$ the$ Newark$ basin$ is$characterized$by$abundant$sporomorphs$of$Triassic$affinity$overlying$the$Lower$Jurassic$CAMP$basalt99.$Note$also$that$all$the$floras$recorded$in$the$Tiefengraben$Member$of$the$Lower$Jurassic$(i.e.$above$the$P.pacificum$beds)$in$the$TJB$stratotype$at$Kuhjoch$(Austria)104$already$existed$in$the$Triassic,$including$C.thiergartii$105.$The$low$δ18O$

New York Canyon (Nevada)

FO P. pacificum

.

McCoy Brook Fm.

North Mtn Basalt

Fundy Basin(Newark SGr)

50 m

1

3

± 0.14 Ma

4

MISSING PART OF RHAETIAN

LO C. crickmayi

FO P. spelae

N13N11

N9

N6cN6b

N3

AC5

1 m -30 -26Corg13δ

Levanto (N Peru)

2

1 m

FO Nevadaphyllites

LO C. crickmayi

FO P. spelaeTJB

JURA

SSIC

201.32 ± 0.13 Ma

201.39 ± 0.14 Ma

201.51 ± 0.15 Ma

LM4-100/101

LM4-90

LM4-86TRIA

SSIC

Rhae

tian

Het

tang

ian

4

10 m

Cwood (PDB)13δ

East Greenland

TJB201.48 ± 0.021 Ma

-1-2 1 20

5

O18δ

Lavernock

FO P. planorbis

TJB

(SW Britain)

10°CVariation in temperature

1 m

?

-28 -26 -24

1000 2000CO2 (ppmV)

1 1

2

2

33

NYC-N10 201.39

-26-27 -25 -24

200.916 ± 0.064 Ma*

201.274 ± 0.032 Ma*

201.52 ± 0.034 Ma*

200 m

6Newark basin

200040006000

pCO2 (ppm)56

Corg13δ

?

?

CAMP


3

values$measured$in$Oysters$coexisting$with$Psiloceras!planorbis$at$the$Lavernock$section106$is$in$agreement$with$super$greenhouse$conditions$pointed$out$by$the$high$pCO2$ inferred$ from$East9Greenland89,90$and$ in$ the$Newark$Basin102$continental$sections.$$Pliensbachian – Toarcian boundary $Relations!between!the!Pl?To!and!Lower!Toarcian!OAE!extinctions!with!the!onset!of!the!Karoo?Ferrar!LIP!!

There$is$a$general$agreement$that$the$activity$of$the$Karoo$Ferrar$LIP$is$likely$related$to$the$crises$that$affected$the$marine$faunas$during$the$Late$Pliensbachian$to$Late$Toarcian$interval.$The$volcanic$pulses$were$short$lived$and$discretely$ distributed$ over$ time1079109,$ which$ is$ consistent$with$ the$ hypothesis$ of$ a$ volcanogenic$ origin$ for$ the$biotic$ crises$ that$ are$ recorded$ by$ the$ ammonites$ during$ this$ period.$ Contact$ metamorphism$ generated$ by$repetitive$intrusions$of$sills$within$the$organic$rich$Ecca$Group$in$the$Karoo$basin$(South$Africa)$was$proposed$as$a$mechanism$to$produce$large$volumes$of$methane,$potentially$inducing$super$greenhouse$conditions,$responsible$for$ the$ Early$ Toarcian$OAE$ and$ the$ negative$ carbon$ isotope$ excursion$ (CIE)110,111.$ High$ resolution$U9Pb$ zircon$geochronology$ from$dolerite$sills$ in$ this$basin$ indicate$a$mean$age$of$182.7$+/9$0.4$Ma$ for$ the$sill$ intrusions112.$This$age$has$recently$been$revised$by$Corfu$et$al.$(ref.$113)$to$183.05$±$0.4$Ma,$based$on$a$recalibration$of$the$Oslo$tracer$solution$against$ the$EARTHTIME$synthetic$100$Ma$solution.$A$new$study$combining$U/Pb$high$precision$dating$on$a$series$of$Karroo$sills$and$on$zircon$from$ash$beds$from$the$lower$Toarcian$Palquilla$sequence$(top$of$the$tenuicostatum$zone,$southern$Peru)$allow$to$correlate$the$onset$of$Karoo$sill$emplacement$with$the$CIE$typical$for$ the$ inferred$ Toarcian$ OAE108.$ Accordingly$ the$ oldest$ sills$ dated$ in$ the$ Karoo$ basin$ are$ 183.014$ ±$0.054/0.072/0.21$Ma$(X/Y/Z$notation$after$Schoene$et$al.;$ref.$114)$while$the$age$of$ash$bed$2$associated$at$the$onset$of$the$CIE$for$the$Toarcian$OAE$in$the$Palquilla$sequence$is$183.22$±$0.25/0.26/0.32$Ma$108(Fig.$2).$Burgess$and$co9authors$(ref.$109)$confirm$the$age$estimate$ for$ the$early$volcanism$in$Karoo$even$the$obtained$age$on$a$granophyre$from$the$New$Amalfi$Sheet$in$southeastern$South$Africa$is$slightly$older$that$the$sills$dated$by$Sell$et$al.$(ref.$108)$–$New$Amalfi$sheet$I9247:$182.246$±$0.045/0.066/0.21$Ma$109.$The$finding$of$the$negative$CIE$in$the$Pacific$area$confirms$its$global$significance1159117.$$The$end9Pliensbachian$extinction,$which$predates$the$Toarcian$OAE$by$probably$several$hundred$kyr$is$marked$by$ an$ important$ diversity$ drop$ associated$ with$ a$ generalized$ sedimentary$ hiatus$ linked$ to$ a$ marked$ marine$regression$ in$NW9Europe$and$ the$Pacific$area118.$As$ indicated$ in$ the$main$ text,$ this$ regression$may$represent$a$major$ short9lived$ glaciation1199124.$ This$ major$ regression$ is$ marked$ by$ important$ emersion$ topography$documented$by$the$deposition$of$thick$shallow$marine$conglomerates$in$the$Dunlap$Formation$in$Nevada$(USA)$containing$the$lower$Toarcian$ammonite$Tiltoniceras$(=$Muller$and$Ferguson’s$(ref.$125)$“Harpoceras”),$and$in$the$top$Pliensbachian$of$the$Ururoa9Kawhia$area,$New$Zealand,$just$below$beds$with$ammonites$that$we$identified$as$the$ coeval$ Dactylioceras$ aff.$ semicelatum$ (Lower$ Toarcian$ in$ age)126.$ The$ regression$ related$ to$ cooling$ is$supported$by$several$recent$δ18O$data$on$belemnites1209122,$127.$$

$PART II. Derivation of a model for sulfur degassing by thermal erosion of the cratonic lithosphere $To$estimate$the$amount$of$SO2/H2S$that$could$be$released$by$thermal$erosion$of$the$cratonic$lithosphere,$we$first$established$a$mass$balance$for$the$amount$of$S$stored$in$the$cratonic$ lithosphere$beneath$Karoo9Ferrar$and$the$CAMP$areas.$For$that,$we$need$(a)$the$volume$of$thermally$erodable$cratonic$lithosphere$and$(b)$the$S$budget$of$the$metasomatized$cratonic$lithosphere.$$Volume of lithosphere thermally eroded !

The$total$surface$area$covered$by$Karoo9Ferrar$and$CAMP$lava$flows$or$Dike$swarms$are$estimated$to$be$~2.15$x$106$km2$(ref.$128$and$references$therein)$and$between$~1$x$107$and$11$x$107$km2$$(refs$129,130),$respectively.$In$our$model,$we$used$the$lower$band$of$CAMP$surface$estimates.$$The$ thickness$ of$ the$ Karoo9Ferrar$ and$ CAMP$ lithosphere$ is$ an$ important$ parameter$ of$ our$ interpretation.$ As$indicated$in$the$article,$different$data$suggest$that$the$Karoo9Ferrar$and$CAMP$have$been$emitted$on$the$top$of$an$initial$thick$lithosphere.$Figure$3$of$the$article$compares$the$position$of$the$peridotite$solidus131$with$the$thickness$


4

of$ the$ lithosphere$ as$ a$ function$ of$ pressure$ and$ temperature.$ Mantle$ melting$ is$ controlled$ by$ the$ rock$ types$present$ in$ the$ upwelling$ plume$ (peridotite/$ pyroxenite)$ and$ their$ respective$ volatile$ contents.$Melting$ of$ such$heterogeneous$ mantle$ in$ an$ upwelling$ plume$ will$ start$ at$ a$ depth$ with$ the$ melting$ of$ volatile9rich$ peridotite$followed$by$pyroxenites$(e.g.$refs.$132,$133).$It$is$important$to$note,$however,$that$the$chemical$composition$of$the$initial$low9degree$melts$produced$from$pyroxenitic$or$volatiles9rich$lithologies$differ$significantly$from$LIP$lavas$134,135.$The$enclosing$peridotite$will$start$to$melt$only$at$relatively$shallow$depth.$Most$models$for$the$formation$of$LIP$ lavas$ suggest$ a$ high$ degree$ of$ partial$melting$ (e.g.$ ref.$ 136)$ in$ order$ to$ explain$ the$ formation$ of$ high9Mg$basalts$ such$ as$ picrites$ or$ komatiites$ observed$ in$ some$ LIPs.$ For$ example,$ Heinonen$ and$ Luttinen$ (ref.$ 137)$suggested$that$Vestfjella$meimechites,$observed$in$Antarctica,$are$produced$at$pressure$between$5$and$6$GPa$at$anomalously$ high$mantle$ potential$ temperatures$ (Tp>1,600°C).$ Hole$ (ref$ 138),$ using$modeled$ primary$magma$composition,$has$estimated$ lower$pressure$ (>$2$GPa)$and$ lower$Tp$ (1450°C$±$50°C)$ for$CAMP$and$Ferrar$LIPs$tholeiitic$ lavas.$ For$ the$ degree$ of$ partial$melting$ relevant$ to$ produce$magma$with$MgO$ content$ higher$ that$ 15$wt.%$ (F$ >5%),$ the$ effect$ of$ volatiles$ on$ melting$ temperature$ is$ insignificant$ as$ indicated$ by$ various$parameterizations$ (e.g.$ refs$139,$ 140).$The$ solidus$ temperature$ for$dry$peridotite$ could$ therefore$be$used$as$ a$minimum$ estimate$ for$ the$ conditions$ of$ magma$ generation$ of$ continental$ flood$ basalt,$ independent$ of$ the$mechanism$ of$ magma$ generation$ in$ a$ mantle$ plume137,$ by$ thermal$ heating$ of$ the$ lithosphere141,142$ or$ by$delamination$of$the$lithosphere143.$Assuming$a$potential$adiabat$temperature$of$1650°C,$the$beginning$of$melting$appears$at$a$depth$of$~$175$km$and$temperature$around$~1700°C.$Taking$into$account$estimates$of$the$thickness$of$ subcontinental$ lithosphere$ beneath$Karoo$ and$ the$ CAMP$ area$ (230$ km$ for$Karoo9Ferrar$ based$ on$ xenoliths$carried$by$Limpopo$Kimberlite144)$and$more$than$200$km$for$the$CAMP$area145,$this$suggests$that$at$least$the$first$25$km$of$subcontinental$lithosphere$needs$to$be$thermally$eroded$to$allow$the$mantle$to$produce$basaltic$lavas.$$$Estimates of potential S stored in the basal part of the lithosphere. !

The$table$SI91$lists$the$different$values$used$to$calculate$the$potential$S$stored$in$the$basal$part$of$the$lithosphere,$Sulfur$that$could$be$potentially$release$to$the$atmosphere$by$thermal$erosion$/$thermal$heating$of$the$lithosphere.$Figure$SI92$reports$the$schematic$composition$of$the$lithosphere$for$the$Karoo$area$proposed$by$Griffin$et$al.$(ref.$144).$The$metasomatic$ lithologies$ (pyroxenite$veins$and$melt9metasomatized$peridotite)$are$ interpreted$as$ the$result$of$the$percolation$of$asthenospheric$melts$and$fluids$associated$with$Archean$and$Proterozoic$subduction$events$ and$ intraplate$ magmatism146.$ One$ key$ aspect$ is$ that$ sulfide$ minerals$ are$ precipitated$ during$ these$metasomatic$event146,147$and$represent$a$S9reservoir$potentially$important$for$the$formation$of$LIPs.$

$Fig. SI- 2. Subcontinental lithosphere section for southern Africa, showing the relative abundances of different peridotite types based on the study of garnet compositions from mantle xenoliths carried by kimberlites144. This section is based on Limpopo kimberlites erupted 500 Ma ago, i.e. before the Karoo-Ferrar volcanic activity.

Schematic section of cratonic lithospheric mantle(modified from Griffin et al., 2003)

Fertile peridotites

Depleted metasomatizedperidotites

100

150

200

0 50 100%

Enclosingmetasomatizedperidotite(S content: 50 - 800 ppm)

Pyroxenite veins(S content: 300-3000 ppm)

Depth

(km)

Melt metasomatizedperidotites

Depletedperidotite(S content:(DMM) 119 ppm)

+


5

For$our$model,$we$assume$ that$ the$basal$part$of$ the$ lithosphere$ is$ composed$of$45%$depleted$peridotite,$ 55%$melt9metasomatised$ peridotite$ and$ 5%$ of$ metasomatic$ veins,$ which$ is$ known$ to$ be$ associated$ with$ the$metasomatic$enrichment$of$the$lithosphere148.$$$Griffin$et!al.$(ref.$146)$indicate$that$the$Sulfur$content$of$peridotite$xenoliths$is$between$50$and$≥6000$ppm.$If$the$extremely$high$S$value$observed$in$one$xenolith$seems$related$to$late$alteration,$the$xenoliths$with$equilibration$temperature$>$1100°C,$ relevant$ for$ the$basal$ part$ of$ the$ lithosphere,$have$>300$ppm$S.$This$ value,$which$ is$ in$agreement$with$that$observed$in$metasomatized$peridotites$studied$by$Jégo$and$Dasgupta$(ref.$149)$or$measured$in$ subcontinental$ peridotitic$ massifs$ (Fig.$ SI93),$ is$ used$ in$ our$ calculation$ as$ the$ S$ content$ of$ metasomatized$peridotite.$A$value$of$119$ppm$is$used$for$ the$sulfur$content$of$depleted$peridotite$according$the$S9estimate$ for$DMM$by$Salters$and$Stracke$(ref.$150).$The$sulfur$content$(800$ppm)$of$metasomatic$veins,$mostly$pyroxenites,$is$derived$from$our$own$compilation$of$pyroxenite$veins$in$orogenic$peridotite$massifs$(Fig.$SI93),$corresponding$to$lithosphere$metasomatized$by$subduction$zone$fluids$or$melts.$ The$ S$ content$was$ calculated$by$using$ the$ concentration$ of$ S$ in$ the$ lithospheric$mantle$ and$ the$ volume$of$ the$thermally$eroded$ lithosphere.$Bockrath$et$al.$ (ref.$151)$have$experimentally$determined$ the$sulfide$solidus$and$liquidus$ in$peridotite$ for$ pressures$ ranging$between$0$ and$3.3$GPa.$Neglecting$ the$ fact$ that$ the$P9T$ range$was$lower$ than$ the$ range$ expected$ for$ thermal$ heating$ of$ the$ lithosphere,$ the$ study$ indicates$ that$ the$ solidus$temperature$ of$ (Fe,Ni,Cu)19x$ monosulfide$ is$ lower$ than$ the$ solidus$ for$ anhydrous$ peridotite.$ Based$ on$ these$experiments,$Hart$and$Gaetani$(ref.$152)$note$that$‘for!a!reasonable!mantle!potential!temperature!(1500°C),!mantle!upwelling!along!an!adiabat!will! intersect! the! sulfide! solidus!at!~160!km!depth,!and! the! sulfide!will!be! fully!molten!

before!the!upwelling!even!reaches!the!peridotite!solidus!at!~110!km’.$This$suggests$that,$during$thermal$erosion$or$thermal$heating$of$lithosphere,$sulfides$are$likely$to$melt.$Nevertheless,$a$S$bearing$phase$seems$to$still$be$present$even$if$the$degree$of$melting$reached$by$the$peridotite$is$significant153,$questioning$the$amount$of$S$release$to$the$melt.$ Though$ solubility$ of$ S$ in$ silicate$ melts$ is$ well$ constrained$ at$ low$ pressure$ (e.g.$ ref.$ 154),$ less$ data$ are$available$at$higher$pressure.$Based$on$a$parameterization$of$Mavrogenes$&$O’Neill$ (ref.$155)$ the$S$ solubility$ in$basaltic$melts$at$a$depth$of$110$km$was$estimated$to$be$around$1000$ppm152.$The$potential$transport$of$S$during$thermal$erosion$of$the$lithosphere$will$be$significantly$enhanced$if$fluids$or$sulfide$liquids$are$produced156.$Recent$high9pressure$experiments$on$sulfide9bearing$ocean9crust$indicate$that,$if$fluids$are$produced$at$depth,$such$fluids$could$dissolve$several$wt.$%$S$largely$independent$on$fO2$149.$$$

$$

Fig. SI-3. Compilation of Sulfur contents in metasomatized peridotites146 compared to peridotites and pyroxenites from subcontinental peridotite massifs. References for peridotite S content are: Ivrea zone – ref. 167; Lherz – ref. 168; Massif Central – refs 169,170; Peridotites Montferrier (southern France) – ref. 171. References for pyroxenites S content: Lherz – ref. 172; Hawaii – ref.; Malenco –ref. 174. * Selected values for peridotite (depleted MORB mantle150), metasomatized peridotite and pyroxenite.

$

0

5

10

15

20

25

30

35

40

0 100 200 300 400 500 600 700 800 900 1000

Subcontinental peridotites Subcontinental pyroxenitesand metasomatized peridotites

Ivrea

Lherz peridotites

Massif Central

Montferrier

Frequency

Bulk rock S content (ppm) Bulk rock S content (ppm)

0

5

10

15

20

25

30

35

40

0 500 1000 1500 2000 2500 3000 3500 4000 4500 5000 5500 6000

Hawaii pyroxenite

Kaapvaal ‘metaso-

matic’ peridotites

Lherz Pyroxenite

Malenco Pyx-ite

* D

ep

lete

d p

erid

otite

(1

19

pp

m)

* M

eta

so

ma

tize

d p

erid

otite

(3

00

pp

m)

* P

yro

xe

nite

s /

Me

taso

ma

tic v

ein

s

(8

00

pp

m)

Frequency


6

We$ suggest$ the$ following$ dynamic$model$ for$ the$ release$ of$ S$ from$ the$metasomatized$ lithosphere.$ This$model$illustrated$in$figure$4$(main$article),$suggests$that$thermal$erosion$of$the$lithosphere$is$linked$to$the$arrival$of$a$thermal$ plume$ or$ by$ internal$ heating$ of$ the$ upper$mantle141,142.$ Note$ that$ our$ hypothesis$ is$ still$ valid$ if$ other$mechanisms$for$CFB$magma$production$are$considered$such$as$delamination$of$the$base$of$the$lithosphere143;$the$thermal$heating$of$the$base$of$the$lithosphere$is$required$in$all$models$to$account$for$the$formation$conditions$of$CFB.$Our$model$ is$ based$ on$ thermal$ erosion$ processes$ such$ as$ those$ described$ in$ the$Ronda$peridotite$massif$(Spain)$by$Van$der$Wal$&$Bodinier$(ref.$157);$Lenoir$et!al.,$(ref.$158)$and$Bodinier$et!al.$(ref.$159),$extrapolated$to$higher$pressure$and$distinct$(metasomatized)$lithospheric$mantle$composition.$During$the$thermal$heating$of$the$base$ of$ the$ lithosphere,$ initial$ melt$ is$ produced$ by$ melting$ of$ phlogopite9bearing$ lithologies$ and$ pyroxenites$present$in$the$lower$part$of$this$lithosphere,$lithologies$characterized$by$solidus$temperatures$significantly$lower$than$ those$ of$ anhydrous$ peridotite.$ The$migration$ of$ these$ volatile$ rich$melts$will$ produce$ km9scale$ pervasive$melt$percolation$associated$with$an$important$accumulation$of$melt$at$the$recrystallisation$front157.$After$several$km$of$ thermal$erosion$of$ the$ lithosphere,$ the$melt$present$at$ the$recrystallization$front$ is$expected$to$be$highly$enriched$in$volatiles$including$sulfur.$We$assume$that$fluids$and/or$sulfide$liquids$form$at$this$stage.$Such$fluids$or$sulfide$melts$are$expected$to$move$to$the$surface$and$provide$a$way$to$emit$significant$S$in$the$form$of$SO2$or$H2S$ to$ the$ atmosphere,$ slightly$ before$ or$ concomitant$ with$ the$ first$ magma$ pulses$ produced$ from$ the$asthenosphere.$ Jégo$&$Dasgupsta$ (ref.$ 149)$ indicate$ that$ at$ high$pressure$ and$ for$ reducing$ conditions,$ the$ S$ is$dissolved$in$fluids$as$H2S.$Nevertheless,$during$the$transport$and$release$to$the$atmosphere,$a$significant$portion$of$ the$H2S$will$be$converted$ to$SO2,$ in$particular$because$ the$oxidation$state$of$ the$ fluids$ is$expected$ to$change$with$decreasing$pressure.$(note:!we!do!not!calculate!the!H2S/SO2!ratio!of!the!gas!flux!to!the!atmosphere,!we!report!the!total!amount!of!S!release!as!SO2$in$our$calculation).$$$

$

Table SI-1. Estimation of the amount of S stored in the basal part of cratonic lithosphere

Volume of lithosphere potentially thermally erodedKaroo-Ferrar CAMP

Surface area covers by continental flood basalt Km2 2150000 10000000

Surface area of lithosphere potentially thermal eroded (=0.5 x Surface covers by CFB) Km2 1075000 5000000Thickness of the lithospheric mantle potentially thermal eroded Km 25 25

Volume of the lithospheric mantle potentially thermally eroded Km3 26875000 125000000

S content in the basal part of the lithosphere: Proportion of the differnt lithologies S content (ppm)

Depleted peridotite 0.4 119 ppmMetasomatized peridotie 0.55 300 ppm

Pyroxenite 0.05 800 ppm

S content: 252.6 ppm

Convert in SO2 505 ppmMass of lithospheric mantle potentially thermally eroded(= volume of lith. x mass of 1 km3 of lithospheric mantle (3.35 x 1012 kg)

9.003 x 1019 kg 4.188 x 1020 kg9.003 x 1016 t 4.188 x 1017 t

Mass of SO2 potentially releasable to the atmosphere(= mass of lithosphere thermally eroded x SO2 content in the lithosphere)

4.547 x 1013 t SO2 2.115 x 1014 t SO2

≈ 45000 Gt SO2 ≈ 210000 Gt SO2

For comparison, the eruption of Pinatubo have release 20 Mt of SO2 in the atmosphere (Bluth et al., ref. 161) while the Laki release 122 Mt of SO2 (Thordarson & Self, ref. 160)


7

If$we$compare$the$amount$of$SO2$calculated$here$with$the$estimated$values$for$the$Laki$(122$MT$SO2;$ref.$$160)$or$Pinatubo$(20$Mt$SO2;$ref.$161)$eruptions,$the$volumes$of$SO2$potentially$realizable$to$the$atmosphere$during$the$initial$stage$of$Karoo$or$CAMP$magmatism$are$large.$As$indicated$in$the$article,$we$assume$that$only$a$fraction$of$the$Sulfur$ initially$stored$ in$ the$basal$part$of$ the$ lithosphere$ is$release$ to$ the$atmosphere,$but$ this$could$create$multiples$injections$of$SO2$at$high$flux$in$the$atmosphere$during$a$relatively$short$period$of$time.$If$the$timing$is$difficult$ to$ constrain,$we$assume$ that$ SO2$ release$ starts$0.2$ to$0.4$Ma$prior$ to$ the$emission$of$CFB$until$massif$magma$ eruptions.$ This$ duration$ is$ based$ on$ the$ thermomechanical$ model$ for$ plume$ lithosphere$ interaction$developed$for$the$case$of$Siberian$Traps162,$model$that$ indicates$a$time$decoupling$of$>0.4$Ma$between$the$first$release$of$gases$from$the$plume$(mostly$CO2$and$HCl)$and$the$first$lavas$eruption.$According$the$different$studies$on$the$climatic$effect$of$volcanic$SO2$degasing$160,161,1639166,$we$hypothesize$that$these$multiple$pulses$could$initiate$cooling$ of$ the$ atmosphere$ and$ global$ biologic$ crises$ recorded$ at$ Rhaetian$ 9$ Hettangian! and$ Pliensbachian$ –$Toarcian$boundaries.$$$$References $$86.$ Guex J. et al. Geochronological constraints on post-extinction recovery of the ammonoids and carbon cycle perturbations during the

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