Geol. Mag.: page 1 of 19 c© Cambridge University Press 2015 1doi:10.1017/S0016756815000126

Modelling the impact of pulsed CAMP volcanism on pCO2 andδ13C across the Triassic–Jurassic transition

AV I V BAC H A N∗‡† & J O NAT H A N L . PAY N E∗∗Department of Geological and Environmental Sciences, Stanford University, Stanford, California, USA

‡Current Address: Department of Geosciences, Pennsylvania State University, Pennsylvania, USA

(Received 2 May 2014; accepted 18 February 2015)

Abstract – A sharp negative δ13C excursion coincides with the end-Triassic mass extinction. This isfollowed by a protracted interval of 13C enrichment. These isotopic events occurred simultaneouslywith the emplacement of the Central Atlantic Magmatic Province (CAMP). Here we use a carbon cyclebox model to explore the effects of episodic carbon release – constrained by recently developed high-resolution chronology – on atmospheric pCO2, ocean chemistry and the δ13C of the ocean–atmospherecarbon pool. Our results are consistent with previous modelling efforts in suggesting that the sharpnegative δ13C excursion and acidification event associated with the extinction are best explained bythe rapid release (<20 ka) of highly 13C-depleted carbon (−70‰). However, our model also indicatesthat the likely short duration of the excursion requires organic carbon burial to have closely followedcarbon injection. The age within the Hettangian of the large positive δ13C excursion which follows iscurrently uncertain. If early Hettangian in age, then our modelling indicates that the interval of 13Cenrichment was closely associated with the volcanic CO2 pulses and pCO2 peaks. If late Hettangian inage, then the 13C enrichment must have lagged the carbon input substantially (by hundreds of thousandsof years) and was associated with CO2 drawdown and over-cooling. Our modelling highlights the needfor improved age constraints on Hettangian stratigraphic sections in order to test between two distinctand contrasting possibilities: continuing carbon cycle instability due to recurrent perturbations fromCAMP activity or a delayed recovery arising from internal biosphere dynamics.

Keywords: end-Triassic mass extinction, Triassic–Jurassic Boundary, carbon isotopes, carbon cycle,geochemical modelling.

1. Introduction

Anthony Hallam is well known for his pioneering workon the end-Triassic extinction. His work was the first tohighlight the association of the biotic crisis with oceananoxia and sea-level changes (Hallam, 1981; Hallam &Wignall, 1999). Perhaps less appreciated, he was alsothe first to collect high-resolution carbon isotopic dataspanning the extinction horizon (Hallam & Goodfel-low, 1990; Hallam, 1994; Morante & Hallam, 1996)and was therefore the first to identify what we nowrecognize as one of the hallmarks of the end-Triassicextinction: the negative δ13C excursion associated withthe disappearance of the Triassic biota.

The stratigraphically abrupt drop in δ13C of carbon-ate and organic carbon coincident with the last occur-rence of Triassic biota has now been identified in nu-merous, globally distributed Triassic–Jurassic bound-ary (T–J) sections (Hallam, 1994; Ward et al. 2001;Hesselbo et al. 2002; Galli et al. 2005; Ruhl, Kür-schner & Krystyn, 2009; Ruhl et al. 2011; Greeneet al. 2012). It has also been the focus of severalmodelling studies aimed at constraining the mass ofCentral Atlantic Magmatic Province (CAMP) -derivedvolatiles and/or volume of methane released (Beerling& Berner, 2002) as well as the associated impacts on

†Author for correspondence: aub33@psu.edu

ocean chemistry such as acidification (Berner & Beer-ling, 2007; Paris et al. 2012) and warming (Huynh& Poulsen, 2005). These models, which focused onthe boundary interval, assumed a complete return tosteady state and full recovery after the initial negativeδ13C excursion. However, recently developed carbonisotope records extending through Hettangian succes-sion have shown that the negative δ13C excursion isfollowed by a protracted interval of 13C enrichment(Fig. 1). This portion of the T–J carbon cycle perturba-tion has yet to be explored by modelling, leaving incom-plete our understanding of the events associated withthe end-Triassic mass extinction and ensuing recov-ery. Moreover, the recently developed high-resolutionchronologies for CAMP activity (Schoene et al. 2010;Blackburn et al. 2013), coupled to pCO2 estimates fromproxy measurements on palaeosols interbedded withthe basalts (Schaller, Wright & Kent, 2011; Schalleret al. 2012), allow the investigation of the relationshipbetween volcanic activity and 13C enrichment (and byinference, ocean chemistry) in unprecedented detail.

Here we use a global biogeochemical box modelto explore the carbon cycle consequences of CAMP-associated perturbations and to constrain the conditionsunder which the recovery from the end-Triassic massextinction may have taken place. In particular, we at-tempt to answer the question of how pCO2 and ocean

2 A . BAC H A N & J. L . PAY N E

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Figure 1. (Colour online) Carbon isotope record of the uppermost Triassic and lowermost Jurassic from the Lombardy Basin, Italy.This is one of the geochemically best-characterized sections presently available with both organic and carbonate carbon showingco-enrichments, an indication that the δ13C signal is primary in origin and representative of the original ocean water chemistry. Thelarge negative excursion at the extinction (N1) is followed by a small positive excursion (P1) and a large protracted excursion (P2).Data from Bachan et al. (2012). δ13Corg hollow circles, δ13Ccarb all other symbols.

carbonate chemistry may have changed during the re-covery and what bounds can be placed on their valuesin light of the δ13C records.

2. Model description

Our model comprises a mass box representing the totalcarbon in the ocean–atmosphere system, an isotopic-mass box representing the overall isotopic compositionof that carbon, as well as boxes representing the totalcalcium and phosphorus in a well-mixed ocean (Fig. 2,Tables 1–3). Inputs to the boxes correspond to deliveryof carbon and alkalinity from volcanism and weatheringto the ocean. Outputs from the boxes correspond to theburial of calcium carbonate and organic matter from theocean in marine sediments (Fig. 2, Table 3). In addition,the carbon box includes a variable Fextra input by whichwe perturb the model.

We set initial model attributes to match a postulatedlatest Triassic steady state (Table 1). The initial valueof dissolved calcium (17 mM) was taken from Horita,Zimmermann & Holland (2002) and the initial valuefor pCO2 (2000 ppm) from Berner (2006). Initial val-

ues for alkalinity (ALK, 1.8 mEq kg−1) and dissolvedinorganic carbon (DIC, 1.9 mM) were chosen to yieldthe appropriate pCO2, together with the ocean calcitesupersaturation commensurate with the modern whole-ocean average of 1.3 (Berner & Beerling, 2007). Weused modern ocean volume (1.32×1021 L) and salinity(35 ppt) to convert concentrations to masses. The val-ues for inputs and outputs from the boxes were takenfrom DePaolo (2004) and Kump & Arthur (1999) andmodified to yield a steady-state δ13C value of +2‰.The baseline magnitude of photosynthetic fractionationwas taken from Hayes, Strauss & Kaufman (1999). Afull list of the model steady-state attributes is given inTable 1.

In the model we calculate ocean chemistry paramet-ers (pH, pCO2 and �) from ALK and DIC at everytime-step using a carbonate system solver based onZeebe & Wolf-Gladrow (2001) and the code in Emer-son & Hedges (2008). The values for ALK and DICare determined from the calcium and carbon boxes,respectively. For simplicity, the contribution of allother conservative ions to ocean alkalinity is assumedconstant.

Modelling the impact of pulsed CAMP volcanism 3

Table 1. Steady-state values for model fluxes and reservoirs. Flux values after DePaolo (2004)and Kump & Arthur (1999). Dissolved calcium based on Horita, Zimmeerman & Holland(2002) and pCO2 after Berner (2006). Magnitude of photosynthetic fractionation from Hayes,Strauss & Kaufman (1999). Concentrations of carbon, calcium and phosphate converted tomasses using an ocean volume of 1.32×1021 L and salinity of 1.035 kg L−1. kALK is the sumof charges from all cations excluding Ca2+.

Description Name Value Unit

Dissolved inorganic carbon DIC 1.9×10−3 mol kg−1

Dissolved calcium Ca 17×10−3 mol kg−1

Alkalinity ALK 1.89×10−3 eq kg−1

Dissolved phosphate PO43− 0.25×10−6 mol kg−1

Constant charge contribution from non-calcium species kALK 32.11×10−3 eq kg−1

Partial pressure of CO2 pCO2 1998 ppmvCalcite saturation � 1.4 –pH pH 7.3 –Steady-state volcanic input Fvolc 16×1012 mol a−1

Weathering input of organic carbon Fworg 8×1012 mol a−1

Weathering input of CaCO Fwcarb 24×1012 mol a−1

Silicate weathering input of calcium Fwsil 12×1012 mol a−1

Weathering input of phosphate Fwp 0.11×1012 mol a−1

Burial flux of organic carbon Fborg 12×1012 mol a−1

Burial flux of CaCO3 Fbcarb 36×1012 mol a−1

Burial flux of phosphate Fbp 0.11×1012 mol a−1

C:P burial ratio CP 106 –δ13C of volcanic flux δvolc –5.5 permilδ13C of carbonate weathering δw

carb 2 permilδ13C of organic carbon weathering δw

org –28 permilδ13C of carbonate burial δ 2 permilδ13C of organic carbon burial δb

org –28 permilPhotosynthetic fractionation ε 30 permil

Table 2. Governing equations

dCdt

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Figure 2. Sketch of model. Ca and P boxes represent the oceanicreservoirs of calcium and phosphate, and the C box representsthe ocean–atmosphere pool of carbon. Inputs and outputs cor-respond to weathering and burial fluxes of carbon, alkalinity andphosphate into and from the ocean. Csys denotes the carbonatesystem solver. Numbers in boxes and adjacent to DIC and ALKare concentrations in mmol kg−1; pCO2 is in units of ppm; andnumbers on fluxes are rates in Tmol (1012 moles) per year. Fluxnames, descriptions and steady-state values given in Tables 1–3.

Table 3. Weathering and burial feedbacks.Fluxes with w superscript representweathering and b burial. Fluxes with isubscript are initial steady-state values. Allothers are time-varying values. Weatheringfeedback parameterization after Berner(2004): G = 3.3, Z = 0.09, Zc = 0.07.

RCO2= CO2

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Feedbacks in the model (Table 3) include a de-pendence of all weathering fluxes on pCO2 using theparameterization of Berner (2004), which includes adifferent weathering coefficient for carbonate rocksreflecting their increased susceptibility to dissolutioncompared to crystalline silicate rocks. Carbonate burialis set to be first-order dependent on the carbonate

4 A . BAC H A N & J. L . PAY N E

saturation state �, such that � > 1 results in carbon-ate burial and � < 1 results in carbonate dissolution(a negative flux). Phosphate burial is set to be first-order dependent on phosphate concentration. Organiccarbon burial is set by the product of phosphate burialand the C:P burial ratio. The C:P burial ratio itself isset to be proportional to phosphate burial, resulting in astrong sensitivity of organic carbon burial to increasedphosphate input.

The model is designed to address a medium range ofdeep-time perturbations: those that are longer than therate of cycling of carbon between surface reservoirs,which is smaller than a few thousand years (ka), butshorter than the hundred million year (Ma) timescaleof fluxes in and out of sedimentary reservoirs (e.g. Gar-rels & Perry, 1974). Consequently, we assume that theflow of carbon between surface reservoirs such as soils,land biota, atmosphere and shallow and deep oceans arein steady state and can be treated as one homogenousreservoir. Because of this assumption of steady state,the model is not suitable for simulating perturbationsoccurring on timescales shorter than the longest in-ternal process, which is the mixing time of the oceans(on the order of a few thousand years; Broecker & Peng,1982). We further leave the treatment of isotopic gradi-ents between the shallow and deep ocean (Kump, 1991)as well as spatial heterogeneities (Panchuk, Holmden& Kump, 2005) for future work. We also assume thatthe imbalance between organic and carbonate carbonweathering and burial is compensated by volcanic in-put, which is itself sourced by metamorphic degass-ing and reverse weathering of the excess burial. Wetherefore assume that the inputs and outputs are closedoutside the model by cycling of carbon between sedi-mentary reservoirs. For readers interested in additionaldetails or running the model on their own, the Matlabcode is provided in the online Supplementary Material(available at http://journals.cambridge.org/geo).

3. Results

Below we present the model response to CO2 injectionsof three different durations: 20, 100 and 500 ka. We ex-plore the effects of differing δ13C values of the injectedcarbon for the 20 ka perturbation, and discuss the im-pacts of changing the model parameters that control thesensitivity of organic carbon burial and attendant 13Cenrichment for the other scenarios. We then assemblethe model runs into a multi-perturbation scenario in aneffort to fit the CAMP degassing chronology and δ13Crecord.

3.a. Single perturbation scenario

The magnitude, duration and isotopic composition ofthe volatile releases from CAMP activity are key vari-ables in modelling the associated carbon cycle perturb-ation. Estimates for the magnitude of volatile releasevary: Schaller et al. (2012) use a value of 3.5×1017 mol(moles) of CO2 to represent a single volatile re-

lease pulse and assume four of them for a total of1.4×1018 mol; Berner & Beerling (2007) use a value of1.76×1018 mol carbon for the total integrated release;and Paris et al. (2012) use, in addition to the value ofBerner & Beerling (2007), a low value of 1.2×1017

and a high value of 2.9×1018 mol for the total volatilerelease. These estimates are generally obtained by mul-tiplying the average volatile load of basalt (0.6 %) bythe reconstructed volume of the province (<4×10 km3;McHone, 2003) times the assumed degassing fraction(100 %).

These estimates are likely conservative however, asthey assume no interaction of the igneous province withthe abundant rift and passive margin sediments throughwhich the CAMP intruded. Geochemical evidencefrom distinctive, fine-grained domains within CAMPvolcanic rocks in the southern Iberian peninsula yieldelevated CaO and 87Sr/86Sr (Martins et al. 2008), sug-gesting a 2–10 % assimilation of a pure CaCO3 end-member which would imply an order of magnitudeincrease in degassed CO2 over unadulterated basalt.Sills in the Amazonian basin intruding into carbonateand evaporite sediments have an average thickness of500 m and cover an area of 1×106 km2 (Marzoli et al.1999). Moreover, CAMP basalts abut coal bodies inthe eastern USA (Robbins, Wilkes & Textoris, 1988),suggesting possible degassing of CO2 and thermogenicmethane from coal (Ogden & Sleep, 2012). The totalmagnitude of interaction between CAMP basalts andsediments is difficult to quantify, but if even a frac-tion of CAMP basalts had a similar magnitude of sed-iment interaction, the amount of carbon released byCAMP may have been substantially underestimated.Similarly, interactions with sediments could lead to aneither more 13C-enriched or more 13C-depleted carbonisotopic signature of the evolved volatiles, dependingupon the isotope composition of the sediments. Volat-iles released from the interaction between basalts andcoals (Retallack & Jahren, 2008; Ogden & Sleep, 2012)would have a δ13C value closer to −20‰ whereas in-teraction with carbonates would lead to a δ13C morepositive than −5.5‰, possibly close to 0‰. A pulseof methane triggered by CAMP volcanism (Beerling &Berner, 2002) would have a value of −50‰ to −100‰depending on whether the gas was thermogenic or bio-genic in origin. Additionally, it is possible (thoughhighly uncertain) that the emitted magmatic carbon it-self may have been 13C-depleted through interactionswith pools of isotopically light carbon in the mantle(Deines, 2002; Cartigny, 2010; Paris et al. 2012).

3.a.1. 20 ka release

The modelled carbon cycle response to a release of3.5×1017 mol of carbon over 20 ka (Fig. 3) is a declinein the δ13C of the ocean (Fig. 3b), a sharp decline fol-lowed by a rise in the carbonate saturation state (�,Fig. 3c) and a 1.5-fold increase in pCO2 (Fig. 3d), fol-lowed by a return to steady state. Assigning more negat-ive δ13C values to the carbon input (Fig. 3b) causes the

Modelling the impact of pulsed CAMP volcanism 5

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Figure 3. Model responses to a single short pulse of carbon simulated by a Gaussian-shaped perturbation with a duration of 20 ka, atotal mass of 3.5×1017 mol, and varying isotopic values. (a) The added carbon flux. (b) δ13C responses for four different carbon isotopicvalues of the input: solid line −5.5‰; dashed line −22‰; dash-dot line −50‰; and dotted line −70‰. (c) Calcite supersaturationresponse. (d) pCO2 response. Total run duration is 200 ka.

minimum δ13C values to decline: with an igneous valueof −5.5‰ (solid line) the δ13C of seawater decreasesfrom a baseline of 2‰ to a minimum of 1.3‰; with aninput of −22‰, corresponding to the input of organiccarbon (dashed line), the δ13C of seawater reaches aminimum of −0.4‰; with an input of −50‰, corres-ponding to thermogenic methane release (dot-dashedline), the δ13C of seawater declines to a minimum of−3.3‰; and with an input of −70‰, correspondingto release of biogenic methane (dotted line), the δ13Cof seawater falls to a minimum of −5.5‰. All minimaare reached c. 13 ka after the start of the perturbation,which nearly coincides with the peak emission rate at10 ka. The corresponding decline in saturation stateis a result of rapidly increasing oceanic DIC, whichreduces the relative abundance of the carbonate ion.A minimum saturation value of 0.99 is reached 11 kaafter the perturbation start. Increased calcium concen-trations due to increased input from weathering andcarbonate dissolution following the CO2 injection leadto a saturation overshoot that peaks at a value of 1.5c. 18 ka after perturbation start.

The pCO2 peak shows two distinct phases character-ized by very different slopes. The first is a sharp peakthat is driven largely by the volatile emission pulse it-self, as is evident by the correspondence between the

two curves. During the perturbation, the input of carbonis rapid enough that inputs of alkalinity from carbonatedissolution and silicate weathering are insufficient tocounter the increase in pCO2. After the perturbationends, the input of alkalinity from silicate weatheringcoupled with burial of carbonate sediments and or-ganic matter drives a gradual reduction in pCO2 backto its pre-excursion baseline. We note that a 20 ka dur-ation for the carbon injection is close to the maximaltimescale to effect whole-ocean acidification, as the re-sponse time of carbonate dissolution is limited to thistimescale by residence time of carbonate in the ocean.The sharp drop in δ13C and carbonate saturation in themodel (Fig. 3) indicates that the perturbation is nearinstantaneous and that releasing carbon on a shortertimescale would make little difference to the modelledresponse.

3.a.2. 100 ka and 500 ka releases

The modelled carbon cycle response to the release of1.5×1018 mol of carbon over 100 ka and the release of3×1018 mol over 500 ka (both with a δ13C of −5.5‰)are given in Figure 4. The main difference betweenthe 20 ka release (modelled above) and the modelledreleases of carbon over hundreds of thousands of years

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Figure 4. Model response to a protracted release of carbon. In all panels solid lines are runs with C:P to the first power and dashedlines are runs with C:P to the third power. Increasing the strength of the feedback on organic carbon burial results in an increase in theδ13C response, a decreased response of the carbonate saturation and a decreased response of pCO2. (a) Excess volcanic input from therelease of 1.5×1018 mol of carbon over 100 ka and 3×1018 released over 500 ka. (b) δ13C responses. (c) Calcite saturation responses.(d) pCO2 responses. Run duration is 800 ka.

is that the latter drive an increase in pCO2 that is of asufficiently long duration to stimulate organic carbonburial and effect substantial 13C enrichments. In thefirst case (1.5×1018 mol over 100 ka), the rate of carbonemission is still rapid enough that a transient declinein δ13C from 2‰ to 1‰ occurs before the burial oforganic carbon increases the δ13C values. In the case of3×1018 mol over 500 ka, the decline in δ13C driven bythe volatile release is completely offset by the increasein δ13C driven by the increase in organic carbon burial.The magnitude of the resulting 13C enrichment in thetwo runs is similar (1.9‰ in the 100 ka run versus2.6‰ in the 500 ka run), indicating that the doublingof the total amount of carbon is approximately offsetby the five-fold increase in the duration over which itis released. This numerical result demonstrates that thetimescales over which the volcanic degassing occurshave an important impact on 13C enrichment. Whileshorter pulses are more effective at driving negativeδ13C excursions, longer releases are required to drivepositive δ13C excursions the longer a release, the moreeffective it is at driving 13C enrichment.

In all cases, the positive δ13C excursion lags peak car-bon input. In the case of the release of 1.5×1018 molover 100 ka, peak δ13C (3.9‰) occurs 78 ka after themaximum rate of carbon emission. In the case of therelease of 3×1018 mol over 500 ka, peak δ13C value(4.6‰) occurs 118 ka after the maximum rate of car-bon emission. Similar trends occur in modelled oceansaturation and pCO2. The release of 1.5×1018 molover 100 ka drives a transient decline in saturation

state (� = 1.2) that is closely associated with peakemissions (only 2 ka later), followed by a satura-tion overshoot (� = 1.7) 24 ka after the maximumin the rate of carbon emissions. In all cases pCO2

reaches a maximum in close association with volat-ile release; in the first case it reaches 4200 ppm only6 ka after peak volatile release and in the secondcase it peaks at 3415 ppm 13 ka after peak volatilerelease.

3.a.3. Sensitivity to anoxia

The inputs and outputs of phosphorus and the degree towhich organic carbon burial is decoupled from P out-puts are key variables in determining the magnitude andtiming of positive excursions in δ13C. Specifically, en-hanced remineralization of phosphate is crucial for pro-ducing and sustaining positive δ13C excursions becauseit allows for a decoupling of the burial of phosphatefrom the burial of alkalinity in a way that weatheringinputs cannot. Volcanic activity can lead to increasedweathering and input of PO4

3− (either by increasingpCO2 and global weathering or by erupting abundant,highly weatherable basalts), leading to increased or-ganic carbon burial associated with the increased phos-phate availability. However, the accompanying increasein the input flux of alkalinity from weathering (whichcan be thought of as the Ca2+ derived from dissolutionof calcium carbonates and calcium silicates) will res-ult in no net change in the fraction of organic carbonburial, and therefore no change in δ13C. For example, a

Modelling the impact of pulsed CAMP volcanism 7

doubling of weathering leading to a doubling of phos-phate input will drive a doubling of organic carbonburial. At the same time however, it will double the in-put of alkalinity and therefore carbonate carbon burialso that, overall, the perturbation to weathering doesnot lead to any steady-state change in δ13C. The mainconclusion is that silicate weathering cannot drive sub-stantial 13C enrichment and that intervals of elevatedδ13C require an additional burial-related feedback onphosphate and organic carbon burial. A corollary isthat if CO2 drawdown is associated with 13C enrich-ment, then silicate weathering must be secondary toorganic carbon burial in effecting the drawdown.

The theoretical argument for the existence of aburial-related phosphate feedback fits well with thelong-standing observation that the burial of phosphateand organic matter can be decoupled via variation inocean redox conditions. Specifically, phosphate is pref-erentially released back into the water column in sed-iments beneath anoxic bottom waters (Ingall, Bustin& Cappellen, 1993; van Cappellen & Ingall, 1994).This mechanism allows large quantities of organic car-bon to be buried without simultaneously sequesteringthe phosphate needed to sustain further productivity orcausing immediate and offsetting increases in the inor-ganic carbon burial flux. Although there are currentlyno records of P accumulation covering the Late Trias-sic – Early Jurassic time interval, observations acrossMesozoic Ocean Anoxic Events (OAEs) – which sharemany traits with the Early Jurassic recovery interval(not the least of which are positive δ13C excursions) –show that both the C:P ratio and absolute concentrationof P increase in tandem with δ13C (Mort et al. 2007;März et al. 2008). The anoxia–productivity feedbacktherefore likely plays a critical role in sustaining theelevated organic carbon burial required for generatingpositive δ13C excursions.

Certain factors can precondition the ocean towardsanoxia and therefore enable the anoxia–productivityfeedback to expand and take hold. Elevated temperat-ures lead to decreased oxygen solubility in the ocean,and certain continental configurations favour nutrienttrapping leading to elevated PO4

3− concentrations incertain parts of the ocean (Meyer & Kump, 2008). Acompilation of PO4

3− and H2S concentrations in mod-ern euxinic basins (fjords, the Black Sea) reveals thatthe two covary and that the linear relation has an in-tercept (interpreted as the threshold for euxinia) at ap-proximately double the modern deep ocean [PO4

3−]concentrations (Meyer & Kump, 2008; Ozaki, Tajima& Tajika, 2011). A 10× increase in phosphate concen-trations leads to pervasive euxinia in a model calibratedto end-Permian conditions (Meyer, Kump & Ridgwell,2008), with peak intensity occurring in oxygen min-imum zones which expand substantially. Sea level canalso play a role in P cycling through its control on thearea of shelf sediments, the chief locus of P and or-ganic carbon deposition. This control is likely to havebeen amplified during times of extensive euxinia on thecontinental shelves.

In our modelling effort, we attempt to capture the es-sence (but obviously not all the details) of the anoxia–productivity feedback (van Cappellen & Ingall, 1994)by allowing the C:P burial of organic carbon to vary asa function of phosphate output, which is itself first-order dependent on the oceanic mass of phosphate(Table 3). When weathering increases as a result of el-evated pCO2, both the modelled inputs and outputs ofphosphate increase. The burial of organic carbon alsoincreases, but in a non-linear fashion relative to thephosphate burial flux (Tsandev, Reed & Slomp, 2012).At steady state, organic carbon burial is tied to P burialvia a C:P ratio of 106:1; this ratio however increaseswhen the P burial flux increases, leading to a strongincrease in organic carbon burial and therefore δ13C.

To explore the sensitivity of the model to the strengthof the parameterized relationship between C:P and Pburial, we triple the exponent on the C:P dependenceon the phosphate burial so that the relationship be-comes cubic instead of linear. Model results under thisassumption are given in Figure 4 in dashed lines. Underthe larger exponent, the δ13C increase is larger (Fig. 4b)and the increases in carbonate saturation and pCO2 aresmaller (Fig. 4c, d) given the same amount of carbonreleased. This behaviour occurs because the strongerthe diversion of the excess volcanic carbon towards or-ganic carbon burial, the smaller the perturbation to thecarbonate-silicate sub-cycle, which is the source of theincreases in � and pCO2. That is, increasing the ‘ease’with which organic carbon burial occurs per increasein pCO2 means that less change in pCO2 is required toachieve the same level of organic carbon burial and 13Cenrichment.

Modifying the starting phosphate concentration inthe ocean also increases the magnitude of the δ13C re-sponse to perturbation but, less obviously, also leadsto a qualitative change in the response of the system(Fig. 5). Increasing the initial phosphate concentrationto 10× and 50× leads to damped oscillations thatpropagate through the entire model, and are expressedin δ13C (Fig. 5b), � (Fig. 5c) and pCO2 (Fig. 5d). Aswe use a variable-step solver tailored for stiff prob-lems (ode15s; Shampine & Reichelt, 1997) the oscil-latory outcome is a genuine model response and notan artefact arising from solver instability or round-offaccumulation.

Oscillations occur when the initial phosphate con-centration is high because the residence time of phos-phate in the model increases as the concentrations in-crease (from c. 3 ka to 150 ka for the extreme, 50×case). In general, in dynamical systems damped oscil-lations (complex eigenvalues of the Jacobian) can occurwhen the relaxation timescales of two state variables aresimilar, and the derivative of one is inversely dependenton the value of other (Zachos & Kump, 2005; Pollard,Kump & Zachos, 2013). In our model, the carbon boxis dependent on the dynamics of P box through theburial of organic carbon, which is P-dependent, whilethe behaviour of the P cycle is dependent on pCO2

and hence the content of the carbon box; oscillations

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500 ka. Dot-dashed line is 1× present oceanic level (0.25 μmol kg−1), dashed line 10× and solid line 50×. Panels as in Figure 4.Increasing the initial phosphate concentration leads to increased residence time for organic carbon burial and results in dampedoscillations in the model. The pCO2 response also shows an undershoot which, in the 50× case, is coincident with peak δ13C. Runduration is 2 Ma.

therefore occur when the response times of the twoboxes become similar.

In this scenario, substantial differences in phase oc-cur between the responses of the different model vari-ables. This phenomenon is most noticeable in the 50×case, where δ13C (Fig. 5b) and pCO2 (Fig. 5d) are ofnearly opposite phase. In this scenario, peak carbon in-put is closely followed by a minimum in δ13C (15 kalater) and a maximum in pCO2 (46 ka later). The sub-sequent peak in δ13C and minimum in pCO2 lag thepeak carbon input substantially (240 ka and 290 kaafter peak input, respectively). Moreover, pCO2 de-clines transiently past its pre-excursion baseline and,due to the lag, coincides with peak δ13C values. Be-cause we do not change the sensitivity of weatheringduring the model run, this intriguing result indicatesthat a pCO2 undershoot can occur even without en-hanced continental weathering and co-occurs with el-evated δ13C.

3.b. Multiple perturbations

Multiple lines of evidence suggest that CAMP erup-tions were recurrent, pulsed and relatively brief innature. Blackburn et al. (2013) used precise U/Pbradiometric dates from zircons in conjunction withastronomically calibrated cyclostratigraphy (Olsen,Schlische & Fedosh, 1996) to constrain the timing oferuption to three main phases: a first major eruptivephase roughly coincident with the main extinction in-terval (HTQ type basalts: Argana Lower, Intermediateand Upper Basalts in Morocco, North Mountain Basaltin the Fundy Basin and Orange Mountain Basalt in the

Newark basin); a second pulse 270 ka later (HFQ/LTQtype basalts: Preakness Basalt in the Newark basin,Hickory Grove basalt in the Culpeper Basin and Holy-oke Basalt in the Hartford Basin); and a third pulse620 ka after the extinction (HFTQ-R: Argana Recur-rent Unit and Hook Mountain Basalt).

The isotopic composition of carbonate soil nodulesdeveloped in palaeosols from sections containing theCAMP basalt flows in North America indicate thatindividual CAMP eruptions strongly impacted atmo-spheric pCO2. The isotopic values indicate that pCO2

increased between two- and three-fold following eachof the three major volcanic episodes (given relative tothe End Triassic Extinction or ETE) at 0, ETE + 270 ka,and ETE + 620 ka (Schaller, Wright & Kent, 2011;Schaller et al. 2012). Additionally, a fourth increase inpCO2 at ETE + 840 does not coincide with any knownvolcanic pulses, though it is well within the overalltime window for CAMP activity (Knight et al. 2004;Nomade et al. 2007; Verati et al. 2007; Marzoli et al.2011). It is important to note that, due to its limitedtime resolution (hundreds of thousands of years), theproxy likely gives a lower time-averaged bound. Theeffect of aliasing may have been particularly importantduring the few hundred years immediately followingeach eruption, when pCO2 may have been substantiallyhigher than indicated by the proxy. The falloff frompeak pCO2 values is observed to occur on timescales ofhundreds of thousands of years which Schaller, Wright& Kent (2011) and Schaller et al. (2012) interpret toresult from silicate weathering. However, c. 90 % of thedecline in pCO2 following a carbon pulse occurs withinthe first 10 ka when carbon dioxide is absorbed by the

Modelling the impact of pulsed CAMP volcanism 9

ocean and removed by carbonate dissolution (Archer,Kheshgi & Reimer, 1997; Rampino & Caldeira, 2011).Only the remaining 10 % of the excess pCO2 is re-moved by silicate weathering on timescales of hun-dreds of thousands of years. We therefore suggest thatthe protracted e-folding times of CO2 drawdown in-dicated by the palaeosol proxy are best interpreted toreflect waning degassing from the basalt flows and/orprotracted thermogenic CO2 devolatilization from in-truded sediments, rather than extremely slow contin-ental weathering following near-instantaneous pulses.In other words, intervals of elevated pCO2 lasting hun-dreds of thousands of years likely require sustainedcarbon input on those same timescales.

Lastly, the δ13C decline immediately below the oldestbasalts in Morocco (Deenen et al. 2010; Dal Corso et al.2014) indicates an earlier injection of 13C-depleted car-bon that preceded the eruption of the first basalts, per-haps by as little as 10–20 ka. Specifically, if the Talcottand Orange Mountain Basalts (above which the pa-laeosol proxy indicates that pCO2 levels increased) areyounger than the Argana basalt, which is underlain bysediments whose δ13C values are already decreasing,then an additional pulse of 13C-depleted carbon predat-ing the basalts is required to accommodate the decliningδ13C values, even in the absence of an indication for anadditional pCO2 pulse in the North American palaeosolrecord. The decline in δ13C below the basalts in Mo-rocco is unlikely to be a result of contact metamorphismfrom the basalts above, as metamorphic devolatiliza-tion would drive off labile components which tend tobe 13C-depleted (Hayes, Kaplan & Wedeking, 1983),leading to more elevated δ13C values in the residuum.The evidence therefore supports a total of five car-bon pulses: an initial short (<20 ka) pulse of highly13C-depleted carbon, and the four releases identifiedby the pCO2 palaeosol proxy (of which the first threecorrespond to known CAMP flows) which we postu-late had a volcanic δ13C signature. A scenario in whichthe first pulse of carbon was related to devolatilizationassociated with intrusion of CAMP basalts into coalsor organic-rich sedimentary basins (perhaps associatedwith establishment of the volcanic plumbing), whilethe other pulses were mostly associated with degassingof the basalt flows themselves, could account nicely forour postulated heterogeneity in the δ13C of the emittedvolatiles.

Assuming a linear relationship between the volcanicforcing and the δ13C response (as in Fig. 4, as opposedto Fig. 5), it is possible to apply five perturbations tothe model with a spacing suggested by the chronologyabove and to tune model and perturbation parametersto the observed δ13C record (Fig. 6). With the timingof the perturbations constrained, there are still threedegrees of freedom left: the duration of the volatile re-lease; the amount of degassed carbon; and its isotopiccomposition. We describe our choices with respect tothe negative N1 excursion (−5‰), the positive P1 ex-cursion (+4‰) and the large protracted P2 excursion(+6‰).

The negative N1 excursion can be readily pro-duced in the model, while simultaneously reprodu-cing the observed doubling of pCO2 from 2000 ppmto 4000 ppm by assuming a 20 ka Gaussian-shapedpulse of 3.5×1017 mol of CO2 with an isotopic valueof −70‰. This value is similar to that calculated byRuhl et al. (2011). The duration (20 ka) of the pulse ofcarbon is sufficiently short to drive a drop in the satur-ation state of the ocean in conjunction with the drop inδ13C.

Replicating the subsequent positive excursion (P1,+4‰) within our model requires a more complexemission scenario. The model indicates that the or-ganic carbon burial generated by a single methane re-lease is too small to drive a +4‰ excursion (Figs 3, 6),and that additional carbon input is required. However, asingle pulse of 3.5×1017 mol C over 60 ka (the inferredmaximal duration of the HTQ basalts) is insufficient todrive such a positive excursion, even assuming com-plete return to a 2‰ baseline following the releaseof the methane (c. 200 ka), a relatively enriched value(−5.5‰) of the emitted carbon and a strong (cubedpower) feedback on organic carbon burial. We find thata Gaussian pulse of 1.2×1018 mol of CO2 released over60 ka, starting when methane peaks, produces the pat-terns of changes observed in the δ13C and pCO2 proxyrecords. We further set the Gaussian to be stronglyleft-skewed so that the combined shape of the two per-turbations (methane and CO2) is a spike that descendssmoothly into a protracted right tail. The reason that wechose this particular parameterization is that, althoughtwo symmetric Gaussian pulses lead to identical effectson δ13C as the single composite, the two pulses lead totwo resolvable peaks in pCO2 and �, phenomena forwhich we currently lack evidence. Following the meth-ane pulse by a more protracted flux of volcanic carbonleads to a rapid recovery and an increase in δ13C, sothat a small positive δ13C excursion results.

Replicating the large (+6‰) P2 excursion requiresan order of magnitude more carbon. To simultaneouslyhonour the carbon isotope record and the pCO2 con-straints, this carbon must be released over a longertimescale. We find that the asymmetric release of2×1018 mol over 150 ka (the inferred duration of theLTQ/HFQ basalts) can account for both observationsquite well. In the model, 3.9-fold and 4.1-fold increasesin organic carbon burial are associated with the P1 andP2 positive δ13C excursions, but both excursions areassociated with only a 1.6-fold increase in carbonatecarbon burial. In absolute terms, peak carbon removalrates are 23 Tmol a−1 (1 Tmol = 1012 mol) via carbon-ate burial but 35 Tmol a−1 and 39 Tmol a−1 via organiccarbon burial. Since carbonate carbon burial repres-ents the sum of carbonate and silicate weathering, thegreater burial of organic carbon (necessary to drive upδ13C values) indicates that the decline in pCO2 fol-lowing the first two peaks is more strongly driven byorganic carbon burial than by silicate weathering.

We set the fourth and fifth volatile releases (atETE + 620 and ETE + 840 ka) to 7.5×1017 of CO2

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at −5.5‰ over 100 ka each in an effort to accountfor the 1–2‰ positive excursions and the approximatedoubling of pCO2 that are observed. Introducing a lar-ger amount of carbon in each pulse results in a largerpCO2 peak (closer to the observed proxy values), butalso produces larger δ13C excursions as well as largerdrops in carbonate saturation, predictions which are in-consistent with the geologic evidence. Even with these‘compromise’ values, we find that a lowering of theanoxia sensitivity in the model is required from cubic(power of 3) to linear (power of 1) to achieve the ap-parently much muted impact of these pulses on δ13C.If the original sensitivity is instead maintained, the res-ulting positive excursions are closer to 2‰ than 1‰in magnitude.

The final perturbation required to account for theisotopic record in the Lombardy Basin is an additionalcarbonate sink. In the Lombardy Basin, the δ13C recordstabilizes at a value close to 1‰ following the P2 ex-cursion, c. 1‰ lower than the 2‰ baseline typical ofthe uppermost Triassic. A similar trend occurs in theQueen Charlotte Islands (Williford et al. 2007), indicat-ing that it is likely a global phenomenon. If so, this shiftin baseline requires an increase in the delivery of alka-linity from weathering relative to carbon from volcan-ism. In our model, we find that increasing weathering by17 % (2×1018 mol increase from the 12×1018 mol a−1

baseline) beginning at 500 ka, together with the de-crease in the anoxia exponent and approximately in linewith the peak of the P2 excursion, can account for thedecline in δ13C baseline values. This increase in weath-ering in the model triggers a decrease in pCO2, as well

as a decline in organic carbon burial. The 17 % increasein weathering we apply is comparable to, though smal-ler than, the 1.5× to 2× increase calculated by Schalleret al. (2012) using the model of Bergman, Lenton &Watson (2004). In this scenario (shift in baseline δ13C),in contrast to the CO2 drawdown associated with posit-ive excursions, the increased delivery of alkalinity fromcarbonate and silicate weathering are the predominantfactors in the CO2 drawdown. It is attractive to specu-late that this model result corresponds to an increase inthe delivery of alkalinity from fresh CAMP basalts thatwere erupted in the tropics, where high temperaturesand abundant rainfall contributed to vigorous chemicalattack and rapid mineral dissolution. Lastly, althoughwe describe the model outputs in terms of the δ13C ofcarbonates, incorporating a variable CO2 and PO4

3−-dependent fractionation (Kump & Arthur, 1999) makesno significant difference in the model output. The δ13Cof organic carbon faithfully tracks the values of thecarbonate carbon δ13C as the response is very nearlysaturated at the high-baseline pCO2 (2000 ppm). Atthis pCO2 the initial fractionation is already very nearthe maximal possible expressed fractionation.

4. Discussion

The key features of the carbon isotopic record of theTriassic–Jurassic transition are the N1, P1 and P2 ex-cursions (Figs 1, 7, 8), which are known from glob-ally distributed sections. The ‘initial’ N1 negative δ13Cexcursion, which co-occurs with the end-Triassic ex-tinction, is one of the most widely observed features

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Modelling the impact of pulsed CAMP volcanism 13

of the T–J carbon isotopic record (Hallam & Goodfel-low, 1990; Hallam, 1994; Morante & Hallam, 1996;McRoberts, Furrer & Jones, 1997; Pálfy et al. 2001;Hesselbo et al. 2002; Guex et al. 2004; Galli et al. 2005;Ward et al. 2007; Ruhl, Kürschner & Krystyn, 2009;Clémence et al. 2010; Deenen et al. 2010; Ruhl, Veld& Kürschner, 2010; Crne et al. 2011; Ruhl et al. 2011;Bachan et al. 2012; Lindström et al. 2012; Dal Corsoet al. 2014). It has even been proposed as a chemostrati-graphic marker for the stage boundary (McRoberts,Ward & Hesselbo, 2007). Although its timing is in-creasingly well constrained (Schaltegger et al. 2008;Schoene et al. 2010), uncertainty remains with respectto its magnitude as recorded in organic and carbonatecarbon (see Greene et al. 2012 and Bachan, van deSchootbrugge & Payne, 2014 for discussion). Never-theless, that a coincidence of extinction, acidificationand a negative δ13C excursion is the null expectationunder a carbon-release scenario (Fig. 3) is a strongargument in favour of a primary exogenic origin forthe N1 excursion. Our modelling joins previous work(Beerling & Berner, 2002; Berner & Beerling, 2007;Paris et al. 2012) in supporting the notion that rapidrelease of large amounts of 13C-depleted carbon can ac-count for both the negative δ13C excursion and a dropin the carbonate saturation state associated with theextinction.

Additionally, our modelling highlights a required as-sociation between the negative carbon-isotope excur-sion and subsequent burial of organic carbon such thata speedy ‘bounce-back’ in δ13C resulted. The durationof the recovery from a release of light carbon, if only bydilution, is between 100 and 500 ka (Fig. 3; Beerling &Berner, 2002). Assuming that the negative δ13C excur-sion represents such a protracted time interval requiresthat a substantial portion of Hettangian time be concen-trated in 1 m of marl in the Lombardy Basin, Italy, whilethe rest of it is represented by over 600 m of sediment; italso requires that in St Audrie’s Bay, UK, up to 500 ka isrepresented in 50 cm of sediment of the Lilstock Form-ation, while the entire Hettangian above (c. 2 Ma) isrepresented by 70 m of Blue Lias Formation (Ruhl et al.2010). In Austria, we find it equally unlikely that somuch time would be concentrated in less than 20 cm ofstrata in the Eiberg Basin in what are otherwise expan-ded sections (Ruhl, Kürschner & Krystyn, 2009). Fur-thermore, the N1 negative excursion in the LombardyBasin is observed to be completely contained withinthe extinction-marl (Bachan, van de Schootbrugge &Payne, 2014). If the extinction-marl is acidification-induced (Hautmann, 2004), its duration, and thereforethat of the excursion, cannot be longer than a few thou-sand years, which is the response time of the carbonatesystem in the ocean (Broecker, 1971). It is thereforeprobable that substantial amounts of organic carbonwere buried immediately following the injection of the13C-depleted carbon such that the isotopic values re-bounded quickly.

Consistent with the supposition that organic car-bon burial immediately followed carbon introduction,

a small (2‰) positive excursion occurs immediatelyabove the N1 negative excursion near the base ofthe Planorbis zone (the P1 excursion). In the UK,this carbon isotope excursion is very well character-ized in both carbonate carbon and organic carbon inLyme Regis, Lavernock Point and Doniford sections(Korte et al. 2009; Clémence et al. 2010). In both theLombardy Basin, southern Alps, Italy (Bachan et al.2012) and New York Canyon, Nevada, USA (Bartoliniet al. 2012), the P1 excursion is distinctly recognizable(Figs 1, 7). It is followed by the much larger positive P2excursion (up to +6‰) recorded in Kennecott Point,Queen Charlotte Islands, British Columbia (Willifordet al. 2007), New York Canyon, Nevada, USA (Bar-tolini et al. 2012) and the Lombardy Basin, southernAlps, Italy (Bachan et al. 2012).

The positive δ13C excursions are likely global andsynchronous. In the Lombardy Basin the 13C enrich-ment associated with P2 excursion occurs in both car-bonate and organic carbon from the same samples(Bachan et al. 2012), a robust indication that it wasalready present in the original water column and is notdiagenetic. The occurrence of the P2 excursion in bothδ13Corg and δ13Ccarb further discounts a compositionalcontrol on δ13C (van de Schootbrugge et al. 2008). Inaddition, the P2 excursion also occurs in organic carbonfrom extra-Tethyan sections in British Columbia andNevada. The P2 excursion is not apparent in the δ13Cof bulk organic carbon in the Newark basin (Whitesideet al. 2010) and only occurs in muted form in SaintAudrie’s Bay (Ruhl et al. 2010; Hüsing et al. 2014).In both sections small-scale variability is much moreprominent than any long-term trend; in Saint Audrie’sBay over 2‰ of bed-to-bed variability occurs, while inthe Newark basin it is closer to 7‰ so that intense localC-cycling may have masked the expression of globalδ13C trends.

A few broader considerations are pertinent to thequestion of synchroneity of the positive excursions. Anincrease in the δ13C of the ocean–atmosphere pool willresult from increased organic carbon burial irrespect-ive of where it occurs: it can be localized and intenseor extra-regional and diffuse. Nonetheless, in placeswhere the rate of carbon export from the surface watersexceeds the rate at which the waters are mixed, local13C enrichment can arise without an accompanying in-crease in organic carbon burial. Some variation in themagnitude of the P2 excursion is apparent, suggest-ing the existence of regional carbon-isotopic gradients:the P2 excursion reaches +6‰ in northern Italy whilereaching only +4‰ in southern Italy (Bachan et al.2012). It is, however, not possible to explain the entiremagnitude of the excursion in this way. The burial of13C-enriched carbonates in one locality (without cor-responding burial of organic carbon) must lead to a re-duction in the δ13C of carbonates deposited elsewheredue to the constraints imposed by mass balance. Infact, the reason that local deviations from the globalδ13C record can exist (e.g. Swart, 2008) is precisely be-cause they are local, and therefore have little leverage

14 A . BAC H A N & J. L . PAY N E

on the global carbon cycle. A positive δ13C excursioncan be local, but it cannot be both ‘local’ and wide-spread at the same time without incurring carbon cycleimbalances.

A historical perspective is also valuable. The ques-tion of whether positive carbon-isotope excursions rep-resent synchronous chronostratigraphic markers hasplayed out numerous times (the Cenomanian–Turonianboundary is a notable early example; Scholle &Arthur, 1980; Bralower, 1988). The debate has so farbeen consistently resolved in favour of the δ13C valuesbeing synchronous and the biostratigraphy diachron-ous. While this is no guarantee that the T–J boundaryδ13C record follows the same trend, it does offer a strongexpectation that it should.

The observed enrichment in 13C indicating elev-ated levels of organic carbon burial following the end-Triassic extinction is also consistent with long-standingobservations that the earliest Jurassic was a period ofwidespread anoxia and deposition of organic-rich sed-iments (Hallam, 1995). For instance, the Hettangian-aged Blue Lias Formation in the UK has TOC levelsin the range of 1–10 % for the majority of the section,even exceeding 10 % in some horizons (Ruhl et al.2010). In the Lombardy Basin, the beds immediatelyabove the extinction marl (belonging to the Malan-otte Formation) are relatively TOC-rich dark-colouredfinely bedded micrites and shales (Galli et al. 2007;Bachan et al. 2012). Nevertheless, studies of organiccarbon burial in modern marine environments (Berner,1982) indicate that the bulk of organic carbon burialoccurs in shelf environments and near river deltas,where abundant clays are present for rapid burial ofadsorbed organic carbon and short oxygen exposuretimes minimize remineralization (Hartnett et al. 1998).In these environments, organic carbon is not highlyconcentrated. Thus, even though black shales are in-dicative of anoxia and locally high productivity, theydo not necessarily constitute an avenue of significantorganic carbon burial. Consequently, though cases ofcorrespondence between the geochemical evidence andsedimentological observations are reassuring, the wide-spread preservation of black shales is not required todemonstrate elevated rates of organic carbon burial.Our modelling, based on the δ13C record, offers a quant-itative estimate of the increase in organic carbon burialwhich was required to produce the positive excursions:in our multi-perturbation modelling scenario 3.9- and4.3-fold increases in organic carbon burial are requiredto produce the P1 and P2 excursions, respectively.

Although the magnitude of the P2 positive δ13Cexcursion is increasingly well-constrained, its timingduring the Hettangian period remains incompletelyunderstood. Careful consideration of the availablebiostratigraphic and radiometric constraints revealsthat there are currently two possible (and conflict-ing) age assignments for the P2 positive excursion,with very different implications for its impacts onocean–atmosphere chemistry. The first assignment is anearly–middle Hettangian (Planorbis–Liassicum) age,

while the other is a middle–late Hettangian (Liassicum–Angulata) age. The first age assignment is based onthe biostratigraphy and isotopic observations fromKennecott Point, Queen Charlotte Islands, BritishColumbia (Carter & Hori, 2005; Ward et al. 2004;Longridge et al. 2007a, b, 2008; Williford et al. 2007)and from the Lombardy Basin, Italy (Kronecker, 1910;Gaetani, 1970; Galli et al. 2007; Jadoul & Galli, 2008;van de Schootbrugge et al. 2008; Bachan et al. 2012).In both these sections, the excursion extends fromnear the Triassic–Jurassic boundary to well below theHettangian–Sinemurian boundary (Fig. 7). In contrastwith the above age assignments (early–middle Het-tangian), Bartolini et al. (2012) assign a late Hettangian(Angulata Zone) age for the P2 excursion, based on theobservation that in New York Canyon, Nevada, the ex-cursion is situated within Angulata-aged beds. Barto-lini et al. (2012) also point out that a large positiveexcursion of the same age occurs in low-resolutiondata in Lyme Regis, UK, found in the supportingmaterial of Whiteside et al. (2010). They further re-interpret the stratigraphy of Longridge et al. (2008)for Kennecott Point so that a late Hettangian age isalso obtained there for the excursion. However, thisreinterpretation implies that the boundary was previ-ously grossly misplaced there (Fig. 8). Available ra-diometric ages highlight the difficulty in reconciling thetwo age assignments. The most recent radiometric agefor the end-Triassic extinction is 201.564 ± 0.015 Ma(Blackburn et al. 2013), a small change from the201.58 ± 0.28 Ma obtained by Schaltegger et al. (2008)and 201.31 ± 0.43 Ma of Schoene et al. (2010), sug-gesting convergence in its absolute age. An ageof 199.43 ± 10 Ma for the Hettangian–Sinemurianboundary was obtained by Guex et al. (2012) froman ash bed in Badouxia-canadensis-containing bedsin Peru, close to the Hettangian–Sinemurian bound-ary (updated from an age of 199.53 ± 0.29 Ma ob-tained by Schaltegger et al. 2008). These data indicatea c. 2 Ma duration for the Hettangian period. Interpol-ating between the tie-points suggests that either the P2excursion spanned slightly less than the first half ofthe Hettangian age (Planorbis–Liassicum zones) andthat significant time passed between its decline and thebase of the Sinemurian (perhaps the entire Angulatazone, up to 1 Ma), or that it began c. 1 Ma into theHettangian age and peaked near its end. The two ageassignments for the P2 excursion are therefore mutuallyexclusive.

Because the record of CAMP pulses in Morocco andNorth America are all dated to within 1 Ma of the ex-tinction and do not extend past the first half of the Het-tangian period (the last CAMP basalt in North Americais dated at ETE + 620 ka, which is at 200.9 Ma; Black-burn et al. 2013), an early Hettangian age for the ex-cursion would suggest synchroneity with the volcanismin the Newark basin and elsewhere (Fig. 7). If, on theother hand, the P2 carbon isotope excursion is late Het-tangian in age, then even the last of the volatile pulses(at ETE + 840 ka) could not have corresponded with

Modelling the impact of pulsed CAMP volcanism 15

it. In that case, the isotope excursion would be coincid-ent with the cooling and pCO2 undershoot identifiedby Schaller et al. (2012) (Fig. 8). Hence, in the firstoption the P1 and P2 positive excursions are associ-ated with warming and pCO2 peaks (Fig. 7), whereasin the second option the P2 excursion is associated withcooling, pCO2 drawdown and even undershoot (Fig. 8).

If the first correlation is correct, then our multiple-perturbation modelling run (Fig. 6) offers a carboncycle scenario that honours both the pCO2 and δ13Cconstraints (Fig. 7). The perturbation scheme com-prises 5 pulses of carbon: the first is highly 13C-depletedpulse (−70‰) which corresponds to the N1 negativeδ13C excursion. It grades smoothly into a second vol-canic carbon emission (−5‰) which corresponds tothe P1 excursion and the third pulse corresponds tothe P2 excursion. The last two pulses have very mutedeffects on δ13C and correspond to CO2 drawdown (pu-tatively due to increased delivery of alkalinity fromCAMP volcanic rocks).

If, however, the latter correlation is correct (Fig. 8),then our model runs in which damped oscillations occuroffer a potential explanation (Fig. 5). In those runs, thepCO2 minimum coincides with peak δ13C values so thatthey are of nearly opposite phase and, most importantly,substantially lag the carbon input pulse. In this scenario,the pCO2 drawdown recognized in the palaeosol proxyrecord would have been driven by enhanced organiccarbon burial, rather than due to increased weatheringof CAMP basalts. Intriguingly, under this correlationthe complex carbon cycle behaviour would be entirelydue to internal dynamics and would be triggered by(but not driven by) the perturbations associated withCAMP volcanism.

If the latter correlation is correct, a key questionarises: if the P2 positive excursion occurs well after thevolcanic carbon pulses have finished, why then does theδ13C record only show a single positive δ13C excursionin the Planorbis zone (the P1 excursion) and no hint ofmultiple pulses (Fig. 8)? The total duration of the fourpCO2 pulses (nearly 1 Ma) is much longer than the res-idence time of carbon in the ocean–atmosphere system,implying that the pulses should be individually distin-guishable in their effects on the global carbon cycle andδ13C. A complete failure of the (global) carbon isotoperecord to record the four individual pulses of carbonor their aliasing somehow into a single positive δ13Cexcursion (the P1 excursion) – on top of the specialconsideration (elevated P concentrations) required toexplain the lag between the carbon pulses and the P2excursion – would constitute a substantial challenge toour understanding of carbon cycle dynamics.

The key to further progress is distinguishing betweenthese very different carbon cycle scenarios by improv-ing the age constraints on the P2 excursion. At leastone prediction that can be offered is that if the P2excursion is late Hettangian as opposed to early Het-tangian in age, the geomagnetic polarity reversals cor-relative to H24R–H27N (Kent & Olsen, 2008; Muttoniet al. 2010) should occur in the Albenza Formation (as

opposed to the Sedrina Limestone) in the LombardyBasin, Italy and in the Sunrise Formation in New YorkCanyon, Nevada.

5. Conclusions

Consistent with previous studies (Beerling & Berner,2002; Berner & Beerling, 2007; Paris et al. 2012), ourmodelling of Triassic–Jurassic carbon cycle dynamicssuggests that the rapid input of highly 13C-depletedcarbon (−70‰) can explain the negative excursion inδ13C and increase in atmospheric pCO2 associated withthe end-Triassic mass extinction (Fig. 3). Such a scen-ario can also account for ocean acidification, consistentwith selective extinction of heavily calcified marine in-vertebrate animals (Kiessling & Simpson, 2011) andlithological evidence for a carbonate dissolution event(Hautmann, 2004). However, our model further indic-ates that rapid rebound from the most negative car-bon isotope values requires that burial of substantialamounts of organic carbon quickly followed the car-bon injection. Such rapid organic carbon burial likelyoccurred primarily on the continental shelves whereexpansion of existing oxygen minimum zones wouldhave been most rapid (on the order of thousands to tensof thousands of years).

Our modelling also indicates that volcanic carbonpulses of sufficient duration and magnitude can giverise to positive δ13C excursions through accelerationof organic carbon burial (Fig. 4). However, when thephosphorus residence time in the model is increaseddamped oscillations occur, indicating that a more com-plex carbon cycle response to a single disturbance ispossible (Fig. 5). In the first case, there is a tight coup-ling between elevated δ13C and CO2 input and the in-creased δ13C is associated with declining (but still el-evated) pCO2 (Fig. 4), whereas in the second case theelevated δ13C is in opposite phase to pCO2 and lags thecarbon input substantially (Fig. 5).

Uncertainty regarding the timing during the Het-tangian period of peak δ13C values (the P2 excur-sion) leaves open two possibilities corresponding tothe two different carbon cycle scenarios. The first op-tion is that the P2 positive excursion is of early–middleHettangian age, which implies that it coincided withCAMP-related volcanic pulses. In this case a multi-pulse model scenario can be fitted to the isotopic re-cord (Fig. 6) and the P1 and P2 excursions corresponddirectly to pulses of carbon from CAMP volcanics(Fig. 7). The second option is that the P2 excursionis of a middle–late Hettangian age, in which case itlagged the CAMP volatile pulses substantially, and wasassociated with pCO2 drawdown and perhaps even un-dershoot (Fig. 8). Our model runs in which dampedoscillations occur offer a potential carbon-cycle-basedexplanation for this observation. If the former scenariois correct, then the δ13C perturbations did not extendpast 500 ka after the extinction, and the biotic recoverywas under way even as the final two pulses of CAMPvolcanism (at ETE + 620 ka and ETE + 840 ka) were

16 A . BAC H A N & J. L . PAY N E

occurring. However, if the latter scenario is correct,then carbon cycle instability persisted for up to a mil-lion years past the cessation of pulsed CAMP activity,and stabilization was only achieved at the beginning ofthe Sinemurian period. The protracted delay betweenthe physical forcing and biotic response in the lattercase implicates ocean biogeochemical dynamics as thecause for the delayed recovery and carbon cycle in-stability. Improving the geochronological constraintson the P2 excursion will help determine whether therecovery from the end-Triassic mass extinction wasdelayed by internal dynamics of the exogenic carboncycle or repeated external perturbations.

Acknowledgements. The authors thank the editor (P. Wig-nall) and an anonymous reviewer for comments which greatlyimproved this manuscript. AB gratefully acknowledges sup-port from a National Science Foundation Graduate ResearchFellowship.

Supplementary material

To view supplementary material for this article, pleasevisit http://dx.doi.org/10.1017/S0016756815000126.

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