Geology

doi: 10.1130/G32473.1 2012;40;79-82Geology

 and Paul B. WignallJacopo Dal Corso, Paolo Mietto, Robert J. Newton, Richard D. Pancost, Nereo Preto, Guido Roghi to the eruption of Wrangellia flood basalts

C spike in the Carnian (Late Triassic) linked13δDiscovery of a major negative   

Email alerting servicesarticles cite this article

to receive free e-mail alerts when newwww.gsapubs.org/cgi/alertsclick

Subscribe to subscribe to Geologywww.gsapubs.org/subscriptions/click

Permission request to contact GSAhttp://www.geosociety.org/pubs/copyrt.htm#gsaclick

official positions of the Society.citizenship, gender, religion, or political viewpoint. Opinions presented in this publication do not reflectpresentation of diverse opinions and positions by scientists worldwide, regardless of their race, includes a reference to the article's full citation. GSA provides this and other forums for thethe abstracts only of their articles on their own or their organization's Web site providing the posting to further education and science. This file may not be posted to any Web site, but authors may postworks and to make unlimited copies of items in GSA's journals for noncommercial use in classrooms requests to GSA, to use a single figure, a single table, and/or a brief paragraph of text in subsequenttheir employment. Individual scientists are hereby granted permission, without fees or further Copyright not claimed on content prepared wholly by U.S. government employees within scope of

Notes

© 2012 Geological Society of America

on September 29, 2014geology.gsapubs.orgDownloaded from on September 29, 2014geology.gsapubs.orgDownloaded from

GEOLOGY, January 2012 79

INTRODUCTIONAfter the end-Permian mass extinction and

the subsequent slow Early Triassic recovery of environments and biota, the Middle–Late Triassic represents a time of relative stability, interrupted only by an episode of more humid conditions known as the Carnian Pluvial Event (CPE) (Simms and Ruffell, 1989). The CPE is characterized by increased input of coarse silici-clastics and a temporary shutdown of carbon-ate systems across the western Tethyan realm (Rigo et al., 2007); high extinction rates of several groups, e.g., ammonoids, crinoids, bryo-zoa (Simms and Ruffell, 1989), and conodonts (Rigo et al., 2007); a fl oral change to more hygrophytic forms (Roghi et al., 2010); and a dramatic increase of conifer resin exudation (Roghi et al., 2006). The CPE occurred at the climax of a long-term increase in global organic carbon burial rates, as suggested by a positive trend in the δ13C record from marine carbon-ate (Korte et al., 2005a) and plant remains (Dal Corso et al., 2011). Compared to the major fi ve mass extinctions (Sepkoski, 1996), the overall Carnian extinction rates are low (Rohde and Muller, 2005), but the CPE coincides with the origin and/or radiation of several important groups (Fig. 1). The fi rst known dinosaurs and calcareous nannoplankton occur in the record shortly after the CPE (Furin et al., 2006), and the radiation of modern conifers was in the Late Triassic (Taylor et al., 2009). What triggered this environmental and biotic turnover remains

unclear. Among the explanations that have been suggested are an increase in rainfall and modifi ed atmospheric circulation in the Tethyan realm driven by the uplift of the Cimmerian oro-gen (Hornung and Brandner, 2005), and a peak expansion and intensity of the Pangean mega-monsoon corresponding to a zenith of continen-tal aggregation (Parrish, 1993). The CPE also coincides with the eruption of the Wrangellia oceanic plateau, suggesting that a causal rela-

tionship might exist (Furin et al., 2006). The onset of large igneous province (LIP) eruptions is often associated with major negative carbon isotope excursions (CIEs) (e.g., Wignall, 2001; Marzoli et al., 2004; Wignall et al., 2009) and mass extinctions, and for at least one LIP, the injection of 12C has been demonstrated to occur at the very onset of basalt eruption (Wignall et al., 2009). However, no carbon isotope pertur-bations have yet been documented for the CPE.

To explore if any carbon cycle perturbation is associated with the CPE, we have performed compound-specifi c isotope analyses (CSIA) of low-molecular-weight (LMW) (n-C17 and n-C19) and high-molecular-weight (HMW) (n-C25 – n-C31) n-alkanes, isoprenoid lipids, and δ13Corg analyses of bulk wood in marine sediments. The δ13C signal of total organic carbon (TOC) was included to improve carbon isotope data resolu-tion around the CPE.

GEOLOGICAL SETTING AND METHODS

Samples were collected from two very well biostratigraphically dated sections of the Dolo-mites (Southern Alps), the Stuores Wiesen and

Geology, January 2012; v. 40; no. 1; p. 79–82; doi:10.1130/G32473.1; 3 fi gures; Data Repository item 2012027.© 2012 Geological Society of America. For permission to copy, contact Copyright Permissions, GSA, or editing@geosociety.org.

Discovery of a major negative δ13C spike in the Carnian (Late Triassic) linked to the eruption of Wrangellia fl ood basaltsJacopo Dal Corso1, Paolo Mietto1, Robert J. Newton2, Richard D. Pancost3, Nereo Preto1, Guido Roghi4, and Paul B. Wignall21Dipartimento di Geoscienze, Università degli Studi di Padova, via Gradenigo 6, 35131 Padua, Italy2School of Earth and Environment, University of Leeds, Woodhouse Lane, Leeds LS2 1JT, UK3 Organic Geochemistry Unit, Bristol Biogeochemistry Research Centre, School of Chemistry, The Cabot Institute, University of Bristol, Bristol BS8 1TS, UK

4Istituto di Geoscienze e Georisorse, CNR, via Gradenigo 6, 35131 Padua, Italy

ABSTRACTMajor climate changes and mass extinctions are associated with carbon isotope anomalies

in the atmosphere-ocean system and have been shown to coincide with the onset of large igne-ous provinces (LIPs) and, by association, their emissions of greenhouse gases and aerosols. However, climatic and biological consequences of some known LIP eruptions have not yet been explored. During the Carnian (Late Triassic) large volumes of fl ood basalts were erupted to form the so-called Wrangellia LIP (western North America). This huge volcanic province is similar in age to a major climatic and biotic change, the Carnian Pluvial Event (CPE), but no evidence of a causal relationship exists other than timing. Here we report a sharp negative δ13C excursion at the onset of the CPE recorded in organic matter. An abrupt carbon isotope excursion of ~−4‰ occurs in terrestrial and marine fossil molecules, whereas total organic carbon records an ~−2‰ shift. We propose that this carbon isotope negative shift was caused by an injection of light carbon into the atmosphere-ocean system linked to the eruption of Wrangellia fl ood basalts. This carbon-cycle perturbation occurs slightly before two major evolutionary innovations: the origin of dinosaurs and calcareous nannoplankton.

Car

nian

TRIA

SS

IC

235

230

Mid

dle

Late

Ladi

nian

Nor

ian

234

233

232

231

229

228

227

226

225

224

236

237

Ma

Wrangellia U-Pb radioisotopic ages

Bio

stra

tigra

phic

ally

-con

stra

ined

age

for W

rang

ellia

eru

ptio

n

Minimum U-Pb radioisotopic age for the CPE

Ishigualasto dinosaurs Ar-Ar radioisotopic age

Firs

t cal

care

ous

nann

opla

nkto

n

Firs

t din

osau

rs

Maximum age range for the CIE

Duration of the CPE

Onset of the CIE

Figure 1. Correlation be-tween Carnian Pluvial Event (CPE) δ13C negative carbon isotope excursion (CIE), Wrangellia eruption, and origin of dinosaurs and calcareous nanno-plankton. Wrangellia U-Pb radioisotopic ages after Greene et al. (2010, and references therein). CPE duration and U-Pb radio-isotopic age after Furin et al. (2006). Ar-Ar radioiso-topic age here reported for Ishigualasto dinosaurs after Furin et al. (2006, and references therein). See the Data Repository (see footnote 1) for a brief discussion of Wrangellia fl ood basalts age. Time scale after Walker and Geissman (2009).

on September 29, 2014geology.gsapubs.orgDownloaded from

80 GEOLOGY, January 2012

the Milieres-Dibona sections (see the GSA Data Repository1 for a map of the study area). Stuores Wiesen is the type section for the base of the Carnian (Mietto et al., 2007) and consists of a series of hemipelagites and thin turbidite beds deposited in a rapidly infi lling basin in hundreds of meters of water depth, adjacent to high relief carbonate platforms. The Milieres-Dibona section is within the Austrotrachyceras austriacum ammonoid zone (latest early Car-nian) and is a succession of marls, limestone, and sandstone deposited in a prodelta-coastal to paralic environment. The upper part of the section records the input of coarse siliciclastics and a switch to a ramp dominated by skeletal carbonates (Preto and Hinnov, 2003). Fossil wood is present as dispersed coalifi ed frag-ments throughout both sections. (For a descrip-tion of methods used for the CSIA and the car-bon isotope analyses of TOC and wood, see the Data Repository.)

RESULTSThe δ13C values of wood fragments (δ13Cwood)

are in the range expected for C3 plants, from −20.1‰ to −25.5‰ (Vienna Peedee belemnite), and do not exceed a range of 2.4‰ within the same bed. Extracted HMW n-alkanes exhibit an odd-over-even carbon number predominance that suggests derivation from higher plant epi-cuticular waxes (Eglinton and Hamilton, 1967) (the carbon preference index, CPI, calculated for n-C25–n-C31 n-alkanes has values from 1.5 to 2.3; see the Data Repository) and suggests that this section remains immature with respect to oil generation. Odd n-C25 to n-C31 n-alkane δ13C values range from −22.0‰ to −30.9‰ and are slightly depleted (1‰–2‰) with respect to bulk wood and TOC (Fig. 2). In marine sediments, n-C17 and n-C19 n-alkanes are usually considered to be mainly of algal origin and thus to repre-sent a marine δ13C signal (Peters et al., 2005): these LMW n-alkanes have δ13C values ranging

from −25.7‰ to −32.5‰. TOC contents vary between 0.43% and 1.34%, and δ13CTOC values range from −21.7‰ to −25.5‰.

A composite record of odd HMW n-alkane, TOC, and wood δ13C values shows a +3‰ long-term increase during the Carnian (Fig. 2) that is consistent with known δ13Cplant data (Dal Corso et al., 2011) and mirrors inorganic car-bon isotope trends derived from brachiopod calcite and marine bulk carbonates (Korte et al., 2005a). This positive shift in marine dis-solved inorganic carbon δ13C was attributed by Korte et al. (2005a) to the redevelopment of coal swamps and an increase in carbon burial after Early Triassic carbon-cycle instability. Within this positive trend, algal and higher-plant n-alkane δ13C and TOC δ13C values reveal a sudden and pronounced negative shift of ~–4‰ and ~–2‰, respectively, in the lower part of the Heiligkreutz Formation corresponding to the lower A. austriacum ammonoid subzone

Figure 2. Composite latest Ladinian–early Carnian δ13C variations of high-molecular-weight n-alkanes, total organic carbon (TOC), and bulk wood. Negative carbon isotope excursion (CIE) occurs at base of Milieres-Dibona section, which corresponds to base of Austrotrachyceras austriacum ammonoid subzone and is characterized by Aulisporites astigmosus palynological assemblage (as defi ned by Roghi et al., 2010). Note positive long-term δ13C shift during Julian substage of Carnian. Regional change in carbonate sedimentation style from high-relief (Cassian) platforms to skeletal ramps after Preto and Hinnov (2003). Palynological assemblages: A–Concentricisporites bianulatus; B—Aulisporites astigmosus; C—Lagenella martinii (Roghi et al., 2010). Pale gray bar shows segment of Milieres-Dibona section containing CIE highlighted in Figure 3.

-30 -28 -26 -24 -21

NO WOOD

0 m

50

Stu

ores

Wie

sen

sect

ion

Mili

eres

-Dib

ona

sect

ion

n-C25n-C27n-C29n-C31

-25 -23 -25 -23 -21-22

Pro

trach

y-ce

ras

Aus

trotra

chyc

eras

aust

riacu

m

Tropitesdilleri

Trac

hyce

ras

aono

ides

Ladi

nian

Car

nian

rego

le-

danu

s D

axat

ina

cf.

cana

dens

is

aon

aus

triac

um

aonoides

oedipus?

nosubzone

Julia

n

Tuva-lian

Chrono-stratigraphy

Ammonoid

Zone

Sub

zone

TRIA

SS

IC

Long

o-ba

rdia

n

Sta

ge

Sub

stag

e

Per

iod

13δ CTOC13δ C wood

13δ C n-alkane

2σ

Sandstone Limestone

Marl and clay

Conglomerate

Dolomites

Legend:

Car

nian

Plu

vial

Eve

nt

Pal

ynol

ogic

alas

sem

blag

es

A

B

C

Ske

leta

l ram

psH

igh

relie

f (C

assi

an) p

latfo

rms

Max

imum

bio

stra

tigra

phic

age

spa

n fo

r Wra

ngel

lia v

olca

nism

1GSA Data Repository item 2012027, map of the study area, description of methods used to perform compound-specifi c isotope analyses, carbon isotope analyses of total organic carbon and wood and palynomorphs extraction, the equation used to calculate the carbon preference index of high-molecular-weight n-alkanes, and a discussion about their source, a brief discussion about the biostratigraphically constrained and radioisotopic age of the Wrangellia large igneous province, and tables containing all carbon isotope data, is available online at www.geosociety.org/pubs/ft2012.htm, or on request from editing@geosociety.org or Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301, USA.

on September 29, 2014geology.gsapubs.orgDownloaded from

GEOLOGY, January 2012 81

and the Aulisporites astigmosus palynological assemblage (as defi ned by Roghi et al., 2010) (latest early Carnian) (Figs. 2 and 3). This CIE is not associated with any change in the depo-sitional environment, as seen from facies in the fi eld. In some samples it was possible to mea-sure the δ13C values of isoprenoid lipids pris-tane and phytane: in marine environments these molecules mostly derive from phytol, which in turn is derived from the phytyl moiety of chlo-rophyll (Didyk et al., 1978); consequently, they are thought to represent a marine phytoplankton C isotope signal. The δ13Cphytane record shows an ~–3‰ shift that parallels that of the n-alkanes and TOC data; the δ13Cpristane data show an ~+5‰ shift corresponding to the fi nal rebound of the negative CIE (Fig. 3).

DISCUSSIONOur carbon isotope data clearly show that at

the onset of the CPE, CO2 enriched in 12C was injected into the exchangeable reservoirs of the active carbon cycle (atmosphere-ocean-land). One possible source for this is CO2 release by the coeval Wrangellia LIP, either by direct injec-tion of magmatic CO2 or by the triggering of additional CO2 release from some other carbon reservoir. Key to this association is the timing of Wrangellia eruption: biostratigraphic data and radioisotopic ages (Fig. 1) constrain Wrangel-lia volcanism to the late Ladinian–late Carnian (Greene et al., 2010; Furin et al., 2006) (see the Data Repository) and together with mag-netostratigraphic data indicate that the eruption of Wrangellia basalts lasted for at least 2 m.y. (Greene et al., 2010).

A volcanic source of light carbon has been invoked for similar CIEs coinciding with mass

extinctions and some oceanic anoxic events (OAEs) (Wignall, 2001). LIP volcanism released large amounts of gas (mostly CO2 and SO2); however, it is generally assumed that the carbon isotopic composition of volcanic CO2 is too heavy (δ13C = −7‰) to explain the observed large and rapid δ13C shifts. The potential con-tribution of volcanic CO2 to the CIE associated with the Wrangellia volcanism can be assessed with a simple mass balance. Greene et al. (2010) estimated the volume of erupted Wrangellia basalts as 140 × 103 km3, although previous esti-mates suggest larger volumes of as much as 1 × 106 km3 (Lassiter et al., 1995). The volume of fl ood basalts erupted could have been much greater because the Wrangellia oceanic plateau was accreted in an orogenic wedge in the Late Jurassic–Early Cretaceous (Greene et al., 2010), and part of the erupted fl ood basalts could have been subducted. Assuming the larger size given by Lassiter et al. (1995) and that the eruption of 1 km3 of modern basalts emits ~5 × 1012 g of car-bon (McCartney et al., 1990), ~5 × 1018 g of car-bon would have been released during the erup-tion of the Wrangellia LIP. Dickens et al. (1995) calculated that to cause a −2‰ to −3‰ δ13C shift in the present-day exchangeable carbon reser-voirs, a release of 2.7−6.8 × 1020 g of carbon is required, far more than the highest estimates of the Wrangellia carbon release. However, studies of other LIP-related basalts and mantle xenoliths suggest that volcanic CO2 δ13C values (−24‰) could have been lower than is generally assumed (−7‰) (Deines, 2002; Hansen, 2006). No δ13C data are available for Wrangellia fl ood basalts, but if associated volcanic CO2 δ13C val-ues were lower, then it is possible that emission of CO2 from Wrangellia volcanism was entirely

responsible for the observed CIE. Additional light CO2 could have been derived from other sources as a positive feedback associated with increasing temperatures due to the emissions of volcanic greenhouse gases. An increase in sea temperatures could have destabilized methane clathrate hydrate reservoirs, and consequent CH4 release and oxidation to CO2 could have introduced additional light CO2 (clathrate CH4 δ13C = −60‰), increasing the magnitude of the negative CIE (e.g., as proposed for the early Toarcian CIE; Kemp et al., 2005).

The injection of CO2 into the ocean-atmo-sphere system could explain the rise of the car-bonate compensation depth (CCD) recorded in the Lagonegro Basin and the crisis of car-bonate systems at the CPE (Rigo et al., 2007). Increasing pCO2, in the absence of deep ocean carbonate sediments in the early Carnian, could easily result in a decrease of pH and ocean acidifi cation because buffering could occur only in platforms, as suggested by Payne et al. (2010) for the end-Permian event. Equally, raised pCO2 may have caused the acceleration of the hydrological cycle observed during the CPE (Simms and Ruffell, 1989; Hornung and Brandner, 2005; Rigo et al., 2007; Roghi et al., 2010) via the greenhouse effect, increasing the nutrient and siliciclastic supply and contributing to the crisis of shallow-water carbonate systems. The diminished carbonate fl ux from platforms would then have caused or have contributed to the rise of the CCD (Rigo et al., 2007).

CONCLUSIONSThe geochemical, paleontological, and

sedimentological features that characterize the CPE shed new lights on its nature and link it to

MARINETERRESTRIALCovered

LimestoneMarl and clay

Detailed stratigraphic log across the CIE

0

10

20

30

40

m

-23 -21 -32 -30 -28 -26

PristanePhytane

-32 -30 -28 -26-31 -29 -27 -25 -23

13 1313δ C TOC δ C isoprenoidδ C n-alkane

-25

n-C25n-C27n-C29n-C31

n-C17n-C19

Figure 3. Close-up of negative carbon isotope excursion (CIE) at Carnian Pluvial Event. Pris-tane, phytane, and low-molecu-lar-weight n-alkanes (n-C17 and n-C19) are thought to represent marine carbon isotope signal, whereas δ13C of high-molecu-lar-weight n-alkanes represents terrestrial signal that refl ects changes in carbon isotope composition of atmosphere.

on September 29, 2014geology.gsapubs.orgDownloaded from

82 GEOLOGY, January 2012

a global perturbation of the carbon cycle trig-gered by a paroxystic volcanic event. Thus, the CPE appears to be yet another example of a sharp negative CIE coinciding with a LIP erup-tion, climate change, and biotic turnover. How-ever, not all LIPs are associated with a nega-tive CIE. LIPs from the Permian to Jurassic (Emeishan, Siberian Traps, Wrangellia, Cen-tral Atlantic Magmatic Province, and Karoo) produced a negative CIE, but others did not (e.g., Ontong Java, Deccan, North Atlantic). A greater availability of methane hydrates in the Permian–Jurassic oceans could explain this pattern. An observation consistent with this hypothesis is that the CPE, Guadalupian, end-Permian, and Toarcian CIEs all occurred after long-term positive carbon isotope trends (Korte et al., 2005a, 2005b; Suan et al., 2010), which suggest the sequestration of organic matter in sediments, perhaps as large methane hydrate reservoirs. Another possibility is that the volcanic CO2 associated with Permian–Jurassic fl ood basalts had suffi ciently light δ13C values to produce substantial negative CIEs in the ocean-atmosphere system.

ACKNOWLEDGMENTSWe thank Alison Kuhl, James M. Williams, and

the Organic Geochemistry Unit (University of Bris-tol) for valuable help with compound-specifi c isotope analyses; Samuel Allshorn for sample preparation at the School of Earth and Environments (University of Leeds); Aurelio Giaretta (Consiglio Nazionale delle Ricerche Padua), Lorenzo Franceschin, and Sandra Boesso for their help at the Dipartimento di Geosci-enze (Padua) laboratories; Andrea Marzoli (Università degli Studi di Padova) for useful discussions; Enrica Marra for separation of wood from sediments; and C.J. Barwell for lipid extraction. We also thank three anonymous reviewers that greatly helped to improve the manuscript. Part of this research was funded by Università degli Studi di Padova–Progetto di Ateneo, CPDA090175/09 (Manuel Rigo).

REFERENCES CITEDDal Corso, J., Preto, N., Kustatscher, E., Mietto, P.,

Roghi, G., and Jenkyns, H.C., 2011, Carbon-isotope variability of Triassic amber, as com-pared with wood and leaves (Southern Alps, Italy): Palaeogeography, Palaeoclimatology, Pa-laeoecology, v. 302, p. 187–193, doi:10.1016/j.palaeo.2011.01.007.

Deines, P., 2002, The carbon isotope geochemis-try of mantle xenoliths: Earth-Science Re-views, v. 58, p. 247–278, doi:10.1016/S0012-8252(02)00064-8.

Dickens, G.R., Neil, O., Jr., Rea, D.K., and Owen, R.M., 1995, Dissociation of oceanic methane hydrate as a cause of the carbon isotope excur-sion at the end of the Paleocene: Paleoceanogra-phy, v. 10, p. 965–971, doi:10.1029/95PA02087.

Didyk, B.M., Simoneit, B.R.T., Brassel, S.C., and Eglinton, G., 1978, Organic geochemical in-dicators of palaeoenvironmental conditions of sedimentation: Nature, v. 272, p. 216–222, doi:10.1038/272216a0.

Eglinton, G., and Hamilton, R.J., 1967, Leaf epicu-ticular waxes: Science, v. 156, p. 1322–1335, doi:10.1126/science.156.3780.1322.

Furin, S., Preto, N., Rigo, M., Roghi, G., Gianolla, P., Crowley, J.L., and Bowring, S.A., 2006, High-precision U-Pb zircon age from the Tri-assic of Italy: Implications for the Triassic time scale and the Carnian origin of calcareous nannoplankton and dinosaurs: Geology, v. 34, p. 1009–1012, doi:10.1130/G22967A.1.

Greene, A.R., Scoates, J.S., Weis, D., Katvala, E.C., Israel, S., and Nixon, G.T., 2010, The archi-tecture of oceanic plateaus revealed by the volcanic stratigraphy of the accreted Wrangel-lia oceanic plateau: Geosphere, v. 6, p. 47–73, doi:10.1130/ges00212.1.

Hansen, H.J., 2006, Stable isotopes of carbon from basaltic rocks and their possible relation to atmospheric isotope excursions: Lithos, v. 92, p. 105–116, doi:10.1016/j.lithos.2006.03.029.

Hornung, T., and Brandner, R., 2005, Biochro-nostratigraphy of the Reingraben Turnover (Hallstatt Facies Belt): Local black shale events controlled by regional tectonics, cli-matic change and plate tectonics: Facies, v. 51, p. 460–479, doi:10.1007/s10347-005-0061-x.

Kemp, D.B., Coe, A.L., Cohen, A.S., and Schwark, L., 2005, Astronomical pacing of methane release in the Early Jurassic Period: Nature, v. 437, p. 396–399, doi:10.1038/nature04037.

Korte, C., Kozur, H.W., and Veizer, J., 2005a, δ13C and δ18O values of Triassic brachiopods and carbonate rocks as proxies for coeval seawater and palaeotemperature: Palaeogeography, Pal-aeoclimatology, Palaeoecology, v. 226, p. 287–306, doi:10.1016/j.palaeo.2005.05.018.

Korte, C., Jasper, T., Kozur, H., and Veizer, J., 2005b, δ18O and δ13C of Permian brachiopods: A record of seawater evolution and continen-tal glaciation: Palaeogeography, Palaeocli-matology, Palaeoecology, v. 224, p. 333–351, doi:10.1016/j.palaeo.2005.03.015.

Lassiter, J.C., DePaolo, D.J., and Mahoney, J.J., 1995, Geochemistry of the Wrangellia fl ood basalt province: Implications for the role of continental and oceanic lithosphere in fl ood basalt genesis: Journal of Petrology, v. 36, p. 983–1009, doi:10.1093/petrology/36.4.983.

Marzoli, A., and 14 others, 2004, Synchrony of the Central Atlantic magmatic province and the Triassic-Jurassic boundary climatic and biotic crisis: Geology, v. 32, p. 973–976, doi:10.1130/G20652.1.

McCartney, K., Huffman, A.R., and Tredoux, M., 1990, A paradigm for endogenous causation of mass extinctions, in Sharpton, V.L., and Ward, P.D., eds., Global catastrophes in Earth His-tory: Geological Society of America Special Paper 247, p. 125–138.

Mietto, P., and 16 others, 2007, A candidate of the Global Boundary Stratotype Section and Point for the base of the Carnian Stage (Upper Trias-sic): GSSP at the base of the canadensis sub-zone (FAD of Daxatina) in the Prati di Stuores/Stuores Wiesen section (Southern Alps, NE Italy): Albertiana, v. 36, p. 78–97.

Parrish, J.T., 1993, Climate of the supercontinent Pangea: Journal of Geology, v. 101, p. 215–233, doi:10.1086/648217.

Payne, J.L., Turchyn, A.V., Paytan, A., DePaolo, D.J., Lehrmann, D.J., Yu, M., and Wei, J., 2010, Calcium isotope constraints on the end-

Permian mass extinction: National Academy of Sciences Proceedings, v. 107, p. 8543–8548, doi:10.1073/pnas.0914065107.

Peters, K.E., Walters, C.C., and Moldowan, J.M., 2005, The biomarker guide (second edition): New York, Cambridge University Press, 1155 p.

Preto, N., and Hinnov, L., 2003, Unraveling the origin of carbonate platform cyclothems in the Upper Triassic Dürrenstein Formation (Dolo-mites, Italy): Journal of Sedimentary Research, v. 73, p. 774–789, doi:10.1306/030503730774.

Rigo, M., Preto, N., Roghi, G., Tateo, F., and Mietto, P., 2007, A rise in the carbonate compensation depth of western Tethys in the Carnian (Late Triassic): Deep-water evidence for the Carnian Pluvial Event: Palaeogeography, Palaeocli-matology, Palaeoecology, v. 246, p. 188–205, doi:10.1016/j.palaeo.2006.09.013.

Roghi, G., Ragazzi, E., and Gianolla, P., 2006, Trias-sic amber of the Southern Alps (Italy): Palaios, v. 21, p. 143–154, doi:10.2110/palo.2005.p05-68.

Roghi, G., Gianolla, P., Minarelli, L., Pilati, C., and Preto, N., 2010, Palynological correlation of Carnian humid pulses throughout western Te-thys: Palaeogeography, Palaeoclimatology, Pa-laeoecology, v. 290, p. 89–106, doi:10.1016/j.palaeo.2009.11.006.

Rohde, R.A., and Muller, R.A., 2005, Cycles in fossil diversity: Nature, v. 434, p. 208–210, doi:10.1038/nature03339.

Sepkoski, J.J., 1996, Patterns of Phanerozoic extinc-tion: A perspective from global data bases, in Walliser, O.H., ed., Global events and event stra-tigraphy in the Phanerozoic: Berlin, Springer, p. 35–51.

Simms, M.J., and Ruffell, A.H., 1989, Synchroneity of climatic change and extinctions in the Late Tri-assic: Geology, v. 17, p. 265–268, doi:10.1130/0091-7613(1989)017<0265:SOCCAE>2.3.CO;2.

Suan, G., Mattioli, E., Pittet, B., Lécuyer, C., Su-chéras-Marx, B., Duarte, L.V., Philippe, M., Reggiani, L., and Martineau, F., 2010, Secu-lar environmental precursors to early Toarcian (Jurassic) extreme climate changes: Earth and Planetary Science Letters, v. 290, p. 448–458, doi:10.1016/j.epsl.2009.12.047.

Taylor, T.N., Taylor, E.L., and Krings, M., 2009, Pa-leobotany. The biology and evolution of fossil plants (second edition): New York, Academic Press, 1230 p.

Walker, J.D., and Geissman, J.W., compilers, 2009, Geologic time scale: Geological Society of America, doi:10.1130/2009.CTS004R2C.

Wignall, P.B., 2001, Large igneous provinces and mass extinctions: Earth-Science Reviews, v. 53, p. 1–33, doi:10.1016/S0012-8252(00)00037-4.

Wignall, P.B., Sun, Y., Bond, D.P.G., Izon, G., New-ton, R.J., Védrine, S., Widdowson, M., Ali, J.R., Lai, X., Jiang, H., Cope, H., and Bottrell, S.H., 2009, Volcanism, mass extinction, and carbon isotope fl uctuations in the middle Perm-ian of China: Science, v. 324, p. 1179–1182, doi:10.1126/science.1171956.

Manuscript received 19 May 2011Revised manuscript received 29 August 2011Manuscript accepted 30 August 2011

Printed in USA

on September 29, 2014geology.gsapubs.orgDownloaded from