A hallmark of the rapid and massive release of carbon during the Palaeocene–Eocene Thermal Maximum is the global negative carbon isotope excursion. The delayed recovery of the carbon isotope excursion, however, indicates that CO2 inputs continued well after the initial rapid onset, although there is no consensus about the source of this secondary carbon. Here we suggest this secondary input might have derived partly from the oxidation of remobilized sedimentary fossil carbon. We measured the biomarker indicators of thermal maturation in shelf records from the US Mid-Atlantic coast, constructed biomarker mixing models to constrain the amount of fossil carbon in US Mid-Atlantic and Tanzania coastal records, estimated the fossil carbon accumulation rate in coastal sediments and determined the range of global CO2 release from fossil carbon reservoirs. This work provides evidence for an order of magnitude increase in fossil carbon delivery to the oceans that began ~10–20 kyr after the event onset and demonstrates that the oxidation of remobilized fossil carbon released between 102 and 104 PgC as CO2 during the body of the Palaeocene–Eocene Thermal Maximum. The estimated mass is sufficient to have sustained the elevated atmospheric CO2 levels required by the prolonged global carbon isotope excursion. Even after considering uncertainties in the sedimentation rates, these results indicate that the enhanced erosion, mobilization and oxidation of ancient sedimentary carbon contributed to the delayed recovery of the climate system for many thousands of years.


Climate events of the Earth’s past reveal the broad array of Earth’s responses to global carbon cycle perturbations and thus provide valuable insights into the potential impacts of future anthropogenic climate change1. The Palaeocene–Eocene Thermal Maximum (PETM) serves as the best-known analogue for anthropogenic climate change due to the significant amount of carbon and geologically rapid rate of release2. During the PETM, an injection of 4,500–10,000 PgC of 13C-depleted CO2 into Earth’s ocean–atmosphere system over <20,000 years resulted in a global warming of 5–9 °C and a global carbon isotope excursion (CIE) of 3–5‰ recorded in terrestrial and marine sediments2,3,4,5,6,7.

Collectively, carbon isotope records of the PETM reveal a rapid 4–20 kyr onset, followed by 70–100 kyr of sustained negative δ13C values, referred to as the CIE body. A subsequent ~50 kyr recovery phase indicated by a return to pre-event δ13C values suggests a period of CO2 drawdown8,9. To reproduce the shape and magnitude of the CIE, models require an additional release of at least 1,480 PgC (if δ13C = −50‰) during the PETM body10. The carbon input has been simulated as a pulsed carbon cycle feedback, continued release from methane hydrate reservoirs or sedimentary carbon oxidation in response to the initial rapid CO2 release6,10,11,12,13.

Today, ancient sedimentary organic carbon (fossil C) reservoirs emit CO2 when sedimentary deposits are exhumed, eroded and transported. During transport, 15–85% of the remobilized fossil C is oxidized to CO2, spanning the range observed in active margins to passive margins with mobile mud belts14,15, which makes fossil C oxidation a potential major source of atmospheric CO2. Globally, PETM records contain evidence of erosion and sediment transport across landscapes, including eroded kaolinite in marine records16, increased fossil C in Bighorn Basin palaeosols17 and coastal sediments from Tanzania18 and Wilson Lake19, and transported terrestrial material that included detrital magnetite20 and fern spores18 in Mid-Atlantic coast records. Global evidence for remobilization of terrestrial and fossil C during the PETM demonstrates the need to assess fossil C oxidation as an additional CO2 source after the PETM onset.

Here we present the first constraints on fossil C accumulation in marginal marine settings. Further, we use these to constrain the amount of CO2 released from fossil C oxidation. Our results indicate that fossil C accumulation in sediments increased by an order of magnitude (between 15- and 50-fold), which resulted in CO2 release on the order of 103 PgC over the PETM body.

CO2 released from increased remobilization and oxidation of fossil C is a sufficient and plausible mechanism to explain sustained elevated atmospheric CO2 levels and produce a prolonged global CIE. This work demonstrates that changes in continental weathering of carbon-rich sedimentary deposits during global warming events may act as a net source of thousands of petagrams of CO2 on 10–100 kyr timescales.

Elevated accumulation of fossil C in coastal sediments

Shallow marine records from the Maryland Coastal Plain present an opportunity to quantify landscape-scale fluxes of remobilized organic matter from a major drainage basin to the marine system. Upper Palaeocene and Lower Eocene sediments from a <150 m palaeowater depth21 were obtained from the Maryland Coastal Plain region of the Salisbury Embayment, specifically from inner shelf sites South Dover Bridge (SDB) and Cambridge–Dorchester Airport (CamDor)22 (Fig. 1). SDB and CamDor are opportune records to assess sediment and organic matter provenance changes at the land–sea interface due to their increase in sedimentation rate from 1.1 to 16.4 cm kyr–1 and increase in terrestrial sediment flux during the PETM19,22. Coupled with previously published records from Tanzania18,20, we assessed changes in fossil C accumulation in coastal environments during the PETM.

Fig. 1: Percent organic carbon and carbonate and organic isotope records plotted against depth at SDB and CamDor inner shelf sites.
Fig. 1

a, Location of SDB and CamDor (CD) cores (modified from Self-Trail et al.48). The Upper Palaeocene Aquia formation is overlain by 13.8 m (CamDor) and 14.9 m (SDB) of Lower Eocene Marlboro Clay, above which is a contact with the Eocene Nanjemoy Formation. b,c, %TOC, carbonate and organic CIEs are presented for the inner-shelf Palaeo–Potomac transect records of SDB (b) and CamDor (c). The grey shaded areas represent the PETM Marlboro Clay, bounded by red dotted lines. The start of the positive δ13Corg enrichment is marked by a green dotted line. Lithology is designated as: c, clay; st, silt; vfs, very fine sand; fs, fine sand.

We confirmed the terrestrial carbon flux to the shelf at SDB by measuring n-alkane indicators of marine/terrestrial provenance, including the terrigenous/aquatic ratio ((n-C27 + n-C29 + n-C31)/(n-C15 + n-C17 + n-C19)) (refs. 23,24,25) and average n-alkane chain length (ACL20–40) (\(\mathop {\sum }\limits_{i = 20}^{40} \frac{{i[C_i]}}{{[C_i]}}\)) (ref. 26). The terrigenous/aquatic ratio increased from 5.8 ± 3.7 before the PETM to 63.2 ± 75.9 during the PETM (Supplementary Fig. 2), which confirms a major switch from marine- to terrestrial-dominated inputs.

The changes in organic matter source altered the shape and size of the CIE recorded by bulk organic matter. Carbonates exhibit a typical negative excursion of 4.0‰ at SDB21 and 5‰ at CamDor27 (Fig. 1). Although organic matter δ13C values similarly declined by 3.5‰ at SDB and 4‰ at CamDor during the PETM onset, at both sites this was followed by a ~6‰ 13C-enrichment during the PETM body (Fig. 1). We estimate 10–20 kyr lapsed between the onset of the PETM and the observed 13C enrichment, based on a PETM linear sedimentation rate (LSR) of 16.4 cm kyr–1 for the Mid-Atlantic coastal plain sites.

The unexpected ~6‰ 13C enrichment in the PETM body cannot be explained by increased primary production rates, which can decrease 12C selectivity28,29,30,31. In a high CO2 world, an enrichment this significant from diminished fractionation during photosynthesis would have required [PO43–] to increase by a factor of four32, but nannoplankton assemblages indicate relatively consistent nutrient levels through the PETM in the Salisbury Embayment21. A shift from algal- to cyanobacterial-dominated primary production33,34,35 was not a driver of the 13C enrichment, as evidenced by the consistent 2α-methylhopane index values ((2α-methylhopane/(2α-methylhopane + C30-hopane))33, which is typically, though not exclusively, used as an indicator of cyanobacteria-derived biomass inputs33,36 in coastal environments33 (Supplementary Fig. 1).

The 13C enrichment is coincident with enhanced inputs of exogenous fossil C, as indicated by biomarker ratios for thermal maturity. When organic matter is heated to temperatures above ~50 °C on geological timescales, less-stable compounds are altered or lost, which increases the proportion of more-stable compounds25,33. Signatures of elevated thermal maturity were unexpected because Cenozoic sediments in the Salisbury Embayment experienced temperatures that were too low to alter biomarker ratios37,38.

Biomarker indicators measured at SDB and CamDor include C31-homohopane-S/(S + R), C32-homohopane-S/(S + R), C27-hopane Ts/(Ts + Tm) (Ts, trisnorneohopanes; Tm, trisnorhopanes) ratios and the norhopane/hopane ratios, all of which increase with thermal maturation25,39,40, and C29-moretane/(moretane + hopane) ratio, which decreases with maturation25,39,41. All the biomarker ratios reach their maximum values in the same core interval as the most positive organic carbon isotope ratio (δ13Corg) values (Fig. 2), which indicates an enhanced transport of 13C-enriched, thermally mature carbon from upland deposits to the coast at both Mid-Atlantic sites. In particular, the upper Cenomanian Raritan Formation, a shallow marine deposit previously hypothesized as a sediment source to the palaeo coasts of Maryland and Virginia during the PETM42,43, has similar isotopic (δ13Corg = −22‰), kerogen and biomarker thermal properties as the fossil endmember (Supplementary Table 2).

Fig. 2: Biomarker thermal maturity ratios at SDB (top) and CamDor (bottom) demonstrate a source change during the PETM body.
Fig. 2

Thermal maturity ratios are plotted alongside the organic isotope record and shaded by timing during the PETM (light grey, pre-PETM; dark grey, PETM onset; black, PETM body at δ13Corg enrichment onset; white, post-PETM). Ts/(Ts + Tm), C31-homohopane (C31HH)-S/(S + R), C32-homohopane (C32HH)-S/(S + R) and norhopane/hopane (NH/H) increase with thermal maturity; C29-moretane/(moretane + hopane) (C29M/H) decreases with thermal maturity. The PETM interval is bounded by red dotted lines. VPDB, Vienna Pee Dee Belemnite.

We applied independent biomarker mixing models to estimate the fraction of organic carbon derived from thermally mature fossil sources (ffossil) in the sediments of SDB, CamDor, and Tanzania cores. Based on the basin thermal history37, we assumed the contemporaneous carbon (marine or terrestrial)44 had immature C31-homohopane-S/(S+R), and C29-moretane/(moretane + hopane) ratios of 0 and 1, respectively (Xbackground). Fossil C endmembers (Xfossil) were designated as the most thermally mature value for each proxy (that is, C31- and C32-homohopane-S/(S + R) = 0.6 and C29-moretane/(moretane + hopane) = 0)25, to estimate the minimum fraction of fossil C (ffossil) required to produce the observed biomarker records (Xmix). Due to the potential for multiple sediment sources to the Mid-Atlantic coast, biomarker ratios obtained from the Raritan Formation were not used. We calculated ffossil at all the depths in SDB, CamDor and Tanzania records (equation (1)), where X represents biomarker ratios and f is the fraction of carbon from background and fossil sources:

$$X_{{\rm{mix}}} = f_{{\rm{fossil}}} \times X_{{\rm{fossil}}} + \left( {1 - f_{{\rm{fossil}}}} \right) \times X_{{\rm{background}}}$$

In all the records, ffossil increased between the PETM onset and PETM body (Supplementary Table 4). There is no evidence for a decreased productivity on the shelf during the PETM21, which implies the increased ffossil probably represented an increased fossil C flux to the sediments.

We calculated the fossil C mass accumulation rates (MARs) at all sites, which demonstrated a 15–50-fold increase in fossil C delivery during the PETM (Fig. 3). We calculated fossil C MAR using LSRs, dry bulk density (ρ), fraction of fossil C (ffossil) and total organic carbon concentration (%TOC) for every depth at all sites (equation (2)). We assumed the actual sedimentation rates were 0.5–2 times the measured LSR, based on previous work that cited an 80% probability that PETM sedimentation rates were within a factor of 2 of the measured LSR45. For SDB and CamDor, we assumed LSRs of 0.55–2.2 cm kyr–1 before the PETM and 8.2–32.8 cm kyr–1 during the PETM. For Tanzania, we assumed LSRs of 0.5–2 cm kyr–1 before the PETM and 3.5–14 cm kyr–1 during the PETM46. We assumed constant ρ values of 2.65 g cm–3.

$${\rm{MAR}} = {\rm{LSR}} \times \rho \times f_{{\rm{fossil}}} \times \frac{{{\rm{TOC}}}}{{100}}$$

Although the increase in ffossil is subdued, the increased sedimentation rate during the PETM results in fossil C accumulation rate increases of 15–27-fold at Tanzania, 15–50-fold at SDB and 15–20-fold at CamDor. Ultimately, this work demonstrates that even small increases in fossil C concentrations in coastal regions may be indicative of significant increases in fossil C transport from land to sea. Although the efficiency of fossil C transport probably varied with changes in PETM conditions, the processes that led to an enhanced remobilization of fossil C also exposed the fossil C to oxidation during transport.

Fig. 3: MARs of fossil C show an increase in fossil C delivery to sediments during the PETM body.
Fig. 3

MARs are based on fossil C mixing models during the PETM body from Tanzania18,20 (αβ/(αβ + ββ) hopanoid isomerization), SDB (C29-M/M + H, C31-HH-S/S + R), CamDor (C29-M/M + H, C31-HH-S/S + R, C32-HH-S/S + R) and the uncertainty around linear sedimentation rates. The x axis represents the MARs of fossil C, and the y axis represents the depth in each section. Leftward bounds on each plot mark the lower-end MAR estimate and rightward bounds mark the upper MAR estimate based on 0.5–2 times the measured sedimentation rate. The grey shaded areas indicate the PETM interval. Data from all the records show an elevated MAR during the PETM body relative to the background delivery of fossil C.

Global CO2 flux estimated from enhanced fossil C oxidation

The elevated signals of fossil C remobilization were a global phenomenon. Direct observations of fossil C in SDB, CamDor, Wilson Lake, Bighorn Basin and Tanzania sediment records17,18,19, coupled with evidence for enhanced weathering, erosion and sediment transport in most coastal marine records, indicate the PETM climate increased the potential for organic matter remobilization and oxidation. Our quantitative constraints on the additional remobilized carbon provide an opportunity to put bounds on the global CO2 release associated with terrestrial C transport during the PETM, and assess whether fossil C oxidation released enough CO2 to extend the duration of the climate event, as suggested previously by numerous researchers12,47,48.

We estimated the global CO2 release from fossil C oxidation using ffossil and TOC measurements from the Mid-Atlantic shelf (SDB and CamDor) and Tanzania with global estimates of LSR changes, total Eocene shelf area and the proportion of fossil C that escapes oxidation in modern settings. We used ffossil and TOC measurements to calculate the concentration of fossil C ([fossil C]) at each depth (equation (3)). Mean fossil C concentrations during the pre-PETM, PETM onset and PETM body were calculated from a composite record of both Mid-Atlantic shelf sites and separately from the one Tanzania record:

$$\left[ {{\rm{Fossil}}\,{\rm{C}}} \right] = \rho \times f_{{\rm{fossil}}} \times \frac{{{\rm{TOC}}}}{{100}}$$

Using a set of assumptions, we calculated the possible range of additional CO2 release during the PETM body on top of the background (pre-PETM) CO2 release. (1) We assumed the global shelf area during the PETM was similar to that of the Eocene (A = 46.9 × 106 km2) (ref. 49). (2) We assumed the pre-PETM shelf sedimentation rates were 1 cm kyr–1, and PETM sedimentation rates varied between 1–20 times the pre-PETM sedimentation rate (LSRpre = 1 cm kyr–1 and LSRPETM = 1–20 cm kyr–1). (3) We assumed a 100 kyr duration of the PETM body (d= 100 kyr) and did not include the PETM onset due to the uncertainty that surrounds this duration. (4) Using the observations from modern less-oxidizing active margins14 to the more-oxidizing mobile mud belts15, we assumed a range of fossil C remineralization (r = 0.15–0.85) for the PETM. Fossil C concentrations from Tanzania were taken to represent a less-oxidizing environment and were coupled with a low (15%) CO2 remineralization rate to generate a lower bound for the increase in CO2 release. The composite fossil C concentrations from the Mid-Atlantic were assumed to represent highly oxidizing mobile mud belt environments43,48 and were coupled with a high (85%) CO2 remineralization rate to generate the upper bound of additional CO2 remineralization rate. The estimated range of total additional CO2 released from fossil C oxidation (CO2 release) during the PETM body falls within the bounds modelled using the above assumptions (equation (4)).

$$\begin{array}{l}\Delta {\rm{CO}}_2\,{\rm{release}} = \\ \left[ {\left( {[{\rm{Fossil}}\,{\rm{C}}_{{\rm{PETM}}\,{\rm{body}}}] \times {\rm{LSR}}_{{\rm{PETM}}}} \right) - \left( {[{\rm{Fossil}}\,{\rm{C}}_{{\rm{pre}}}] \times {\rm{LSR}}_{{\rm{pre}}}} \right)} \right] \times \frac{{r \times A \times d}}{{\left( {1 - r} \right)}}\end{array}$$

Using these constraints and the proportion of fossil C estimated from the biomarker mixing models (which represents fossil C that escaped oxidation), we estimated that the total additional CO2 release from enhanced fossil C oxidation fell by between 102 and 104 PgC over a ~100 kyr event (Fig. 4). If global average fossil C oxidation during the PETM was between the modern bounds (15%-85%) and global sedimentation rates increased during the PETM similar to coastal observations (~5–15 times increase in LSR), CO2 release was likely on the order of 103 PgC. CO2 budgets during sediment transport will be improved by future studies that better constrain CO2 drawdown from changes in terrestrial biosphere C burial. Using available data and modelled total CO2 release scenarios from fossil C oxidation over the PETM body, we suggest fossil C oxidation was a major contributing factor to the CO2 release that extended the duration of the PETM (Fig. 4).

Fig. 4: Global estimates of fossil-C-derived CO2 release during the PETM body.
Fig. 4

The global estimate of additional CO2 release during the PETM body is displayed as a grey area bounded between two black lines. The lower bound represents the total CO2 release if all the PETM marine margins were less-oxidizing environments or active margins with a high fossil C preservation. The lower bound is calculated from 15% fossil C oxidation rates and is based on observational records from Tanzania. The upper bound represents the global CO2 release if all the marine margins were mobile mud belt passive margins with a low fossil C preservation. The upper bound is calculated from 85% fossil C oxidation rates and is based on composite records from the Mid-Atlantic coastal plain. The x axis represents the sedimentation increase during the PETM relative to pre-PETM conditions. The left y axis designates the total CO2 release in terms of PgC.

The magnitude of the sustained release of CO2 from the weathering of fossil C, some 10–20 kyr after the PETM onset, suggests this was a plausible mechanism to account for the sustained CIE9,10. Further, the global CO2 addition estimated from the model bounds the amount, timing and rates of carbon release required to sustain the deep ocean carbonate dissolution, as presented in previous modelling studies9,10, and are consistent with widely observed landscape-scale signals of sediment and refractory carbon remobilization during the PETM17,18,19,43. Partial oxidation of fossil C stocks in remobilized sedimentary deposits alone or in conjunction with thermogenic methane11 could have released enough CO2 to sustain the negative CIE during the PETM body.

Hyperthermal events trigger long-term CO2 release

This work demonstrates the possibility of long-term (~104–105 yr) carbon-release feedbacks associated with an increased sediment remobilization in high CO2, high temperature environments. Our results indicate an increased remobilization and oxidation of fossil C from sedimentary carbon stocks could have released thousands of petagrams of carbon as CO2 to the ocean–atmosphere system after the PETM onset, in amounts sufficient to explain the sustained high CO2 within the PETM body. Additionally, our results bolster previous hypotheses that high organic matter concentrations in many marine PETM records may be a signature of enhanced carbon transport from land, not solely of enhanced primary productivity19.

Warming and high CO2 of the PETM, coupled with intensification of the hydrological cycle in coastal regions48, promoted an enhanced weathering erosion and sediment transport processes that exposed and oxidized carbon previously locked in sedimentary deposits. The range in our model results underlines the outstanding uncertainty as to climate–landscape–deposition feedbacks with consequences for fossil C exposure, transport, oxidation and burial. For example, modelled timescales of landscape response to changes in precipitation can vary significantly50, and it remains difficult to predict how changes in water and sediment supply affect sediment transport dynamics and deposition within terrestrial and shallow marine environments. Additionally, although carbon remobilization can be a CO2 source when fossil C is oxidized, transport systems also remove CO2 when biosphere carbon is buried and removed from the active carbon cycle. This study does not address how the PETM perturbations might have shifted the relative balance of these processes, but it does demonstrate that the intensity and timescales on which erosion and transport systems both release and drawdown CO2 are critical to our understanding of carbon cycling during the PETM.

The total quantity of carbon released during the PETM is comparable to the predicted anthropogenic release from fossil fuel burning. Although the rates of CO2 release during the PETM are a fraction of the anthropogenic rates, the PETM is considered the best geological analogue for future carbon release8,9. This study highlights the possibility of delayed feedbacks associated with sediment erosion that have the potential to extend the duration of anthropogenic CO2 release. Carbon cycle feedbacks observed during the PETM reinforce the importance of preventing CO2 from reaching dangerous thresholds in the modern, as rapid warming today may trigger irreversible responses that extend the duration of high CO2 and global warming for many thousands of years.


Organic δ13C measurements

Sediment samples were powdered via a ball mill, and 1.5 g of powdered sediment were decarbonated using 15 ml of 1 N HCl for 30 min to 2 h to remove carbonate. Samples were neutralized via repeated rinsing with MilliQ water until the pH rose above 5. %TOC and bulk δ13Corg were measured using a Costech ECS 4010 Elemental Analyzer coupled to a Thermo Delta Plus isotope ratio mass spectrometer. The combustion reactor (length 45.4 cm, diameter 18 mm) was packed with chromium(iii) oxide and silvered cobalt(ii,iii) and heated to 650 °C, and the oxidation reactor (length 45.4 cm, diameter 14 mm) was packed with reduced copper wires and heated to 1,020 °C. The water trap (length 11 cm, diameter 8 mm) was packed with magnesium perchlorate. CO2 and N2 were separated on a gas chromatography column heated to 50 °C prior to transfer to the isotope ratio mass spectrometer. Samples were weighed into tin capsules and analysed with a suite of blanks and five laboratory standards with known isotope and %TOC values (Supplementary Table 4). δ13Corg values and %TOC of the samples were calculated using a three-point standard curve, and δ13Corg values are reported in standard delta notation against Vienna Pee Dee Belemnite. Instrument precision and accuracy were ±0.18‰ and ±0.61‰, respectively. Duplicate samples had a reproducibility of ±0.2‰ for δ13Corg and ±0.05% for %TOC.

Biomarker analyses

Total lipid extracts were obtained via the accelerated solvent extraction of approximately 15 g of sediment and dried to <1 ml under a stream of N2. Lipid extracts were separated into fractions (aliphatic, aromatic and polar) using gravity column separations with a fine mesh silica stationary phase and mobile phases of 100% hexane (aliphatic fraction), 90% hexane 10% methylene chloride (aromatic fraction) and 70% methylene chloride:30% MeOH (v:v) (polar fraction). Aliphatic fractions were analysed via gas chromatography–mass spectrometry on a Thermo Scientific Trace 1310 Gas Chromatograph coupled to a thermo Scientific ISQ LT single quadrupole mass spectrometer. A Restek Rxi-5HT fused silica column (30 m length × 0.25 mm internal diameter × 0.25 µm film thickness) was used with a helium carrier gas. The oven programme started with an injection temperature of 40 °C held for 1.5 min, followed by a rise of 15.0 °C min–1 until the temperature reached 140 °C, after which the temperature increased at 6.0 °C min–1 until it reached 320 °C, at which it was held for 20 min, for a total run time of 58 min. The flow was split and 90% was transferred to an ISQ quadrupole mass spectrometer with a transfer line temperature of 340 °C and electron ionization of 300 °C, which was scanned over the mass range 50–550 amu at 5 scans s–1. The other 10% of the split flow was routed to a flame ionization detector held at 330 °C. Biomarker quantifications were determined from flame ionization detector peak area, using AGSO standard oil, Macondo oil and an n-C10 to n-C40 alkane standard and response curve. Mass spectra were used for compound identification by comparison with standards and published spectra.

Code availability

The mixing model computer code is available from the authors upon reasonable request.

Data availability

The authors declare that the data supporting the findings of this study are available within the paper and its Supplementary Information files. Additional source data for Fig. 4 is available upon reasonable request.

Additional information

Publisher’s note: Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.


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Funding for this study was provided by National Science Foundation grant no. CE-1416663. We acknowledge discussions with M. Robinson, K. Hantsoo and J. A. Grey. We are grateful to D. Walizer for analytical support. Any use of trade, firm or product names is for descriptive purposes only and does not imply endorsement by the US Government.

Author information


  1. Department of Geosciences, Pennsylvania State University, University Park, PA, USA

    • Shelby L. Lyons
    • , Allison A. Baczynski
    • , Timothy J. Bralower
    • , Elizabeth A. Hajek
    • , Lee R. Kump
    • , Ellen G. Polites
    •  & Katherine H. Freeman
  2. Earth & Planetary Sciences Department, University of California, Santa Cruz, CA, USA

    • Tali L. Babila
    •  & James C. Zachos
  3. Eastern Geology and Paleoclimate Science Center, US Geological Survey, Reston, VA, USA

    • Jean M. Self-Trail
  4. Department of Geological Sciences, University of Delaware, Newark, DE, USA

    • Sheila M. Trampush
  5. School of Geosciences, University of Louisiana at Lafayette, Lafayette, LA, USA

    • Jamie R. Vornlocher


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S.L.L., A.A.B., E.G.P. and J.R.V. carried out the organic isotope analyses, S.L.L. and A.A.B. carried out the biomarker analyses, S.L.L. interpreted the biomarker and isotope data, T.L.B. and J.C.Z. conducted and interpreted the carbonate isotope analyses, J.M.S.-T. determined the sedimentation rates, S.L.L. and E.A.H. designed the mixing model, K.H.F., J.C.Z., S.M.T., L.R.K., T.J.B. and A.A.B. contributed to major improvements within the models and data interpretation, S.L.L. wrote the paper, and all the authors contributed to interpreting the data and editing the paper. K.H.F. advised the direction of the research.

Competing interests

The authors declare no competing interests.

Corresponding author

Correspondence to Shelby L. Lyons.

Supplementary Information

  1. Supplementary Information

    Supplementary Methods, Tables and Figures.

  2. Supplementary Data

    Supplementary Data Set.

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