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Articles https://doi.org/10.1038/s41561-018-0277-3

Palaeocene– Thermal Maximum prolonged by carbon oxidation

Shelby L. Lyons 1*, Allison A. Baczynski1, Tali L. Babila2, Timothy J. Bralower1, Elizabeth A. Hajek 1, Lee R. Kump1, Ellen G. Polites1, Jean M. Self-Trail3, Sheila M. Trampush4, Jamie R. Vornlocher5, James C. Zachos2 and Katherine H. Freeman 1

A hallmark of the rapid and massive release of carbon during the Palaeocene–Eocene Thermal Maximum is the global negative carbon 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 bio- marker 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 accumula- tion rate in coastal 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 2 4 and demonstrates that the oxidation of remobilized fossil carbon released between 10 and 10 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 system for many thousands of years.

limate events of the Earth’s reveal the broad array of fossil C in Bighorn Basin palaeosols17 and coastal sediments from Earth’s responses to global perturbations and Tanzania18 and Wilson Lake19, and transported terrestrial material thus provide valuable insights into the potential impacts of that included detrital magnetite20 and fern spores18 in Mid-Atlantic C 1 future anthropogenic . The Palaeocene–Eocene coast records. Global evidence for remobilization of terrestrial and Thermal Maximum (PETM) serves as the best-known analogue fossil C during the PETM demonstrates the need to assess fossil for anthropogenic climate change due to the significant amount of C oxidation as an additional CO2 source after the PETM onset. carbon and geologically rapid rate of release2. During the PETM, Here we present the first constraints on fossil C accumulation 13 an injection of 4,500–10,000 PgC of C-depleted CO2 into Earth’s in marginal marine settings. Further, we use these to constrain the ocean– system over <​20,000 years resulted in a global amount of CO2 released from fossil C oxidation. Our results indi- warming of 5–9 °C and a global carbon isotope excursion (CIE) of cate that fossil C accumulation in sediments increased by an order 2–7 3–5‰ recorded in terrestrial and marine sediments . of magnitude (between 15- and 50-fold), which resulted in CO2 Collectively, carbon isotope records of the PETM reveal a rapid release on the order of 103 PgC over the PETM body. 13 4–20 kyr onset, followed by 70–100 kyr of sustained negative δ​ C CO2 released from increased remobilization and oxidation of values, referred to as the CIE body. A subsequent ~50 kyr recov- fossil C is a sufficient and plausible mechanism to explain sustained 13 ery phase indicated by a return to pre-event δ​ C values suggests a elevated atmospheric CO2 levels and produce a prolonged global 8,9 period of CO2 drawdown . To reproduce the shape and magnitude CIE. This work demonstrates that changes in continental weath- of the CIE, models require an additional release of at least 1,480 PgC ering of carbon-rich sedimentary deposits during global warming 13 10 (if δ​ C =​ −​50‰) during the PETM body . The carbon input has events may act as a net source of thousands of petagrams of CO2 on been simulated as a pulsed carbon cycle feedback, continued release 10–100 kyr timescales. from hydrate reservoirs or sedimentary carbon oxidation 6,10–13 in response to the initial rapid CO2 release . Elevated accumulation of fossil C in coastal sediments Today, ancient sedimentary organic carbon (fossil C) reservoirs Shallow marine records from the Maryland Coastal Plain present emit CO2 when sedimentary deposits are exhumed, eroded and an opportunity to quantify landscape-scale fluxes of remobilized transported. During transport, 15–85% of the remobilized fossil organic matter from a major drainage basin to the marine system. C is oxidized to CO2, spanning the range observed in active margins to Upper Palaeocene and Lower Eocene sediments from a <​150 m passive margins with mobile mud belts14,15, which makes fossil C oxi- palaeowater depth21 were obtained from the Maryland Coastal dation a potential major source of atmospheric CO2. Globally, PETM Plain region of the Salisbury Embayment, specifically from inner records contain evidence of erosion and transport across shelf sites South Dover Bridge (SDB) and Cambridge–Dorchester landscapes, including eroded kaolinite in marine records16, increased Airport (CamDor)22 (Fig. 1). SDB and CamDor are opportune

1Department of Geosciences, Pennsylvania State University, University Park, PA, USA. 2Earth & Planetary Sciences Department, University of California, Santa Cruz, CA, USA. 3Eastern and Paleoclimate Science Center, US Geological Survey, Reston, VA, USA. 4Department of Geological Sciences, University of Delaware, Newark, DE, USA. 5School of Geosciences, University of Louisiana at Lafayette, Lafayette, LA, USA. *e-mail: [email protected]

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a 76° W 74° W

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Fig. 1 | Percent organic carbon and carbonate and organic isotope records plotted against depth at SDB and CamDor inner shelf sites. 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 13 red dotted lines. The start of the positive δ​ Corg enrichment is marked by a green dotted line. Lithology is designated as: c, clay; st, silt; vfs, very fine sand; fs, fine sand. records to assess sediment and organic matter changes during the PETM19,22. Coupled with previously published records at the land–sea interface due to their increase in sedimentation rate from Tanzania18,20, we assessed changes in fossil C accumulation in from 1.1 to 16.4 cm kyr–1 and increase in terrestrial sediment flux coastal environments during the PETM.

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210 cstvfs –30 –25 –20 0 0.5 1 0 0.2 0.4 0.6 0.6 0.4 0.2 0 123 13C (‰ VPDB) δ organic Ts/(Ts + Tm) C31HH-S/(S + R) C29M/H NH/H 210

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Fig. 2 | Biomarker thermal maturity ratios at SDB (top) and CamDor (bottom) demonstrate a source change during the PETM body. 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, 13 PETM body at δ​ Corg 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 confirmed the terrestrial carbon flux to the shelf at SDB by the consistent 2α-methylhopane​ index values ((2α-methylhopane/​ 33 measuring n-alkane indicators of marine/terrestrial provenance, (2α-methylhopane​ +​ C30-hopane)) , which is typically, though not including the terrigenous/aquatic ratio ((n-C27 +​ n-C29 +​ n-C31)/ exclusively, used as an indicator of -derived biomass 23–25 33,36 33 (n-C15 +​ n-C17 +​ n-C19)) (refs. ) and average n-alkane chain inputs in coastal environments (Supplementary Fig. 1). 40 iC[]i 26 13 length (ACL20–40) (∑ ) (ref. ). The terrigenous/aquatic ratio The C enrichment is coincident with enhanced inputs of exog- i=20 []Ci increased from 5.8 ±​ 3.7 before the PETM to 63.2 ±​ 75.9 during the enous fossil C, as indicated by biomarker ratios for thermal matu- PETM (Supplementary Fig. 2), which confirms a major switch from rity. When organic matter is heated to above ~50 °C marine- to terrestrial-dominated inputs. on geological timescales, less-stable compounds are altered or lost, The changes in organic matter source altered the shape and size which increases the proportion of more-stable compounds25,33. of the CIE recorded by bulk organic matter. Carbonates exhibit a Signatures of elevated thermal maturity were unexpected because typical negative excursion of 4.0‰ at SDB21 and 5‰ at CamDor27 Cenozoic sediments in the Salisbury Embayment experienced tem- (Fig. 1). Although organic matter δ​13C values similarly declined by peratures that were too low to alter biomarker ratios37,38. 3.5‰ at SDB and 4‰ at CamDor during the PETM onset, at both Biomarker indicators measured at SDB and CamDor include 13 sites this was followed by a ~6‰ C-enrichment during the PETM C31-homohopane-S/(S +​ R), C32-homohopane-S/(S +​ R), C27-hopane body (Fig. 1). We estimate 10–20 kyr lapsed between the onset of Ts/(Ts +​ Tm) (Ts, trisnorneohopanes; Tm, trisnorhopanes) ratios and the PETM and the observed 13C enrichment, based on a PETM lin- the norhopane/hopane ratios, all of which increase with thermal –1 25,39,40 ear sedimentation rate (LSR) of 16.4 cm kyr for the Mid-Atlantic maturation , and C29-moretane/(moretane +​ hopane) ratio, coastal plain sites. which decreases with maturation25,39,41. All the biomarker ratios The unexpected ~6‰ 13C enrichment in the PETM body can- reach their maximum values in the same core interval as the most 13 not be explained by increased primary production rates, which positive organic carbon isotope ratio (δ​ Corg) values (Fig. 2), which 12 28–31 13 can decrease C selectivity . In a high CO2 world, an enrich- indicates an enhanced transport of C-enriched, thermally mature ment this significant from diminished fractionation during pho- carbon from upland deposits to the coast at both Mid-Atlantic sites. 3– tosynthesis would have required [PO4 ] to increase by a factor of In particular, the upper Cenomanian Raritan Formation, a shallow four32, but nannoplankton assemblages indicate relatively consistent marine deposit previously hypothesized as a sediment source to the nutrient levels through the PETM in the Salisbury Embayment21. palaeo coasts of Maryland and Virginia during the PETM42,43, has 13 A shift from algal- to cyanobacterial-dominated primary produc- similar isotopic (δ​ Corg =​ −​22‰), kerogen and biomarker thermal tion33–35 was not a driver of the 13C enrichment, as evidenced by properties as the fossil endmember (Supplementary 2).

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Tanzania SDB SDB CamDor CamDor CamDor

αβ/(αβ + ββ) C29M/H C31HH-S/(S + R) C29M/H C31HH-S/(S + R) C32HH-S/(S + R) –10 –185 –185 –210 –210 –210

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–20 –195 –195 –220 –220 –220 –25 –200 –200 Depth (m) –225 –225 –225 –30 –205 –205

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0 0 0 0 0 0 0.1 0.2 0.3 0.4 0.5 0.1 0.2 0.3 0.4 0.5 0.1 0.2 0.3 0.4 0.5 0.1 0.2 0.3 0.4 0.5 0.1 0.2 0.3 0.4 0.5 0.1 0.2 0.3 0.4 0.5 Mass accumulation rate of fossil C (gC cm–2 kyr–1)

Fig. 3 | MARs of fossil C show an increase in fossil C delivery to sediments during the PETM body. MARs are based on fossil C mixing models during the 18,20 PETM body from Tanzania (α​β​/(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.

We applied independent biomarker mixing models to estimate Although the increase in ffossil is subdued, the increased sedimen- the fraction of organic carbon derived from thermally mature fos- tation rate during the PETM results in fossil C accumulation rate sil sources (ffossil) in the sediments of SDB, CamDor, and Tanzania increases of 15–27-fold at Tanzania, 15–50-fold at SDB and 15–20- cores. Based on the basin thermal history37, we assumed the con- fold at CamDor. Ultimately, this work demonstrates that even small 44 temporaneous carbon (marine or terrestrial) had immature C31- increases in fossil C concentrations in coastal regions may be indic- homohopane-S/(S + R), and C29-moretane/(moretane +​ hopane) ative of significant increases in fossil C transport from land to sea. ratios of 0 and 1, respectively (Xbackground). Fossil C endmembers Although the efficiency of fossil C transport probably varied with (Xfossil) were designated as the most thermally mature value for each changes in PETM conditions, the processes that led to an enhanced (that is, C31- and C32-homohopane-S/(S +​ R) =​ 0.6 and C29- remobilization of fossil C also exposed the fossil C to oxidation dur- moretane/(moretane +​ hopane) =​ 0)25, to estimate the minimum ing transport. fraction of fossil C (ffossil) required to produce the observed bio- marker records (Xmix). Due to the potential for multiple sediment Global CO2 flux estimated from enhanced fossil C oxidation sources to the Mid-Atlantic coast, biomarker ratios obtained from The elevated signals of fossil C remobilization were a global phe- the Raritan Formation were not used. We calculated ffossil at all the nomenon. Direct observations of fossil C in SDB, CamDor, Wilson depths in SDB, CamDor and Tanzania records (equation (1)), where Lake, Bighorn Basin and Tanzania sediment records17–19, coupled X represents biomarker ratios and f is the fraction of carbon from with evidence for enhanced , erosion and sediment background and fossil sources: transport in most coastal marine records, indicate the PETM cli- mate increased the potential for organic matter remobilization and oxidation. Our quantitative constraints on the additional remobi- Xf Xf(1 ) X mix =×fossil fossil +−fossil × background (1) lized carbon provide an opportunity to put bounds on the global

CO2 release associated with terrestrial C transport during the In all the records, ffossil increased between the PETM onset and PETM, and assess whether fossil C oxidation released enough CO2 PETM body (Supplementary Table 4). There is no evidence for a to extend the duration of the climate event, as suggested previously decreased productivity on the shelf during the PETM21, which by numerous researchers12,47,48. implies the increased ffossil probably represented an increased fossil We estimated the global CO2 release from fossil C oxidation using C flux to the sediments. ffossil and TOC measurements from the Mid-Atlantic shelf (SDB and We calculated the fossil C mass accumulation rates (MARs) CamDor) and Tanzania with global estimates of LSR changes, total at all sites, which demonstrated a 15–50-fold increase in fossil C Eocene shelf area and the proportion of fossil C that escapes oxi- delivery during the PETM (Fig. 3). We calculated fossil C MAR dation in modern settings. We used ffossil and TOC measurements using LSRs, dry bulk density (ρ), fraction of fossil C (ffossil) and total to calculate the concentration of fossil C ([fossil C]) at each depth organic carbon concentration (%TOC) for every depth at all sites (equation (3)). Mean fossil C concentrations during the pre-PETM, (equation (2)). We assumed the actual sedimentation rates were PETM onset and PETM body were calculated from a composite 0.5–2 times the measured LSR, based on previous work that cited an record of both Mid-Atlantic shelf sites and separately from the one 80% probability that PETM sedimentation rates were within a fac- Tanzania record: tor 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 TOC [FossilC] =×ρ ffossil × (3) before the PETM and 3.5–14 cm kyr–1 during the PETM46. We 100 assumed constant ρ values of 2.65 g cm–3. Using a set of assumptions, we calculated the possible range

TOC of additional CO2 release during the PETM body on top of the MARL=×SR ρ ××ffossil (2) 100 background (pre-PETM) CO2 release. (1) We assumed the global shelf area during the PETM was similar to that of the Eocene

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6 2 49 (A =​ 46.9 ×​ 10 km ) (ref. ). (2) We assumed the pre-PETM shelf 105 Percent of fossil C oxidized to CO –1 sedimentation rates were 1 cm kyr , and PETM sedimentation 85 Passive margins rates varied between 1–20 times the pre-PETM sedimentation rate 4 –1 –1 10 (LSRpre =​ 1 cm kyr and LSRPETM =​ 1–20 cm kyr ). (3) We assumed a 100 kyr duration of the PETM body (d = 100 kyr) and did not include 15

release (PgC) 3 the PETM onset due to the uncertainty that surrounds this dura- 2 10 tion. (4) Using the observations from modern less-oxidizing active Active margins 14 15 margins to the more-oxidizing mobile mud belts , we assumed 102 a range of fossil C remineralization (r =​ 0.15–0.85) for the PETM. Additional CO

Fossil C concentrations from Tanzania were taken to represent a 2 101 less-oxidizing environment and were coupled with a low (15%) CO2 0510 15 20 remineralization rate to generate a lower bound for the increase in Relative sedimentation rate increase CO2 release. The composite fossil C concentrations from the Mid-

Atlantic were assumed to represent highly oxidizing mobile mud Fig. 4 | Global estimates of fossil-C-derived CO2 release during the 43,48 belt environments and were coupled with a high (85%) CO2 rem- PETM body. The global estimate of additional CO2 release during the ineralization rate to generate the upper bound of additional CO2 PETM body is displayed as a grey area bounded between two black remineralization rate. The estimated range of total additional CO2 lines. The lower bound represents the total CO2 release if all the PETM released from fossil C oxidation (∆CO2 release) during the PETM marine margins were less-oxidizing environments or active margins body falls within the bounds modelled using the above assumptions with a high fossil C preservation. The lower bound is calculated from (equation (4)). 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 ΔCO2 release= fossil C preservation. The upper bound is calculated from 85% fossil C oxidation rates and is based on composite records from the Mid- [([FossilCPETM body]L× SRPETM) (4) Atlantic coastal plain. The x axis represents the sedimentation increase rA××d during the PETM relative to pre-PETM conditions. The left y axis −×([FossilCprep]LSR re)] × (1−r) designates the total CO2 release in terms of PgC.

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 concentrations in many marine PETM records may be a signature release from enhanced fossil C oxidation fell by between 102 and of enhanced carbon transport from land, not solely of enhanced 104 PgC over a ~100 kyr event (Fig. 4). If global average fossil C oxi- primary productivity19. dation during the PETM was between the modern bounds (15%- Warming and high CO2 of the PETM, coupled with intensifi- 85%) and global sedimentation rates increased during the PETM cation of the hydrological cycle in coastal regions48, promoted an similar to coastal observations (~5–15 times increase in LSR), CO2 enhanced weathering erosion and sediment transport processes 3 release was likely on the order of 10 PgC. CO2 budgets during that exposed and oxidized carbon previously locked in sedimentary sediment transport will be improved by future studies that better deposits. The range in our model results underlines the outstand- constrain CO2 drawdown from changes in terrestrial C ing uncertainty as to climate–landscape–deposition feedbacks burial. Using available data and modelled total CO2 release scenar- with consequences for fossil C exposure, transport, oxidation and ios from fossil C oxidation over the PETM body, we suggest fossil burial. For example, modelled timescales of landscape response to 50 C oxidation was a major contributing factor to the CO2 release that changes in can vary significantly , and it remains dif- extended the duration of the PETM (Fig. 4). ficult to predict how changes in and sediment supply affect

The magnitude of the sustained release of CO2 from the weath- sediment transport dynamics and deposition within terrestrial ering of fossil C, some 10–20 kyr after the PETM onset, suggests and shallow marine environments. Additionally, although car- 9,10 this was a plausible mechanism to account for the sustained CIE . bon remobilization can be a CO2 source when fossil C is oxidized, Further, the global CO2 addition estimated from the model bounds transport systems also remove CO2 when biosphere carbon is bur- the amount, timing and rates of carbon release required to sus- ied and removed from the active carbon cycle. This study does tain the deep ocean carbonate dissolution, as presented in previ- not address how the PETM perturbations might have shifted the ous modelling studies9,10, and are consistent with widely observed relative balance of these processes, but it does demonstrate that the landscape-scale signals of sediment and refractory carbon remobi- intensity and timescales on which erosion and transport systems 17–19,43 lization during the PETM . Partial oxidation of fossil C stocks both release and drawdown CO2 are critical to our understanding in remobilized sedimentary deposits alone or in conjunction with of carbon cycling during the PETM. 11 thermogenic methane could have released enough CO2 to sustain The total quantity of carbon released during the PETM is the negative CIE during the PETM body. comparable to the predicted anthropogenic release from fossil

fuel burning. Although the rates of CO2 release during the PETM Hyperthermal events trigger long-term CO2 release are a fraction of the anthropogenic rates, the PETM is consid- This work demonstrates the possibility of long-term (~104–105 yr) ered the best geological analogue for future carbon release8,9. carbon-release feedbacks associated with an increased sediment This study highlights the possibility of delayed feedbacks associ- remobilization in high CO2, high environments. Our ated with sediment erosion that have the potential to extend the results indicate an increased remobilization and oxidation of fos- duration of anthropogenic CO2 release. Carbon cycle feedbacks sil C from sedimentary carbon stocks could have released thou- observed during the PETM reinforce the importance of prevent- sands of petagrams of carbon as CO2 to the ocean–atmosphere ing CO2 from reaching dangerous thresholds in the modern, as system after the PETM onset, in amounts sufficient to explain rapid warming today may trigger irreversible responses that the sustained high CO2 within the PETM body. Additionally, extend the duration of high CO2 and global warming for many our results bolster previous hypotheses that high organic matter thousands of years.

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Acknowledgements the authors contributed to interpreting the data and editing the paper. K.H.F. advised the Funding for this study was provided by National Science Foundation grant no. direction of the research. 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 Competing interests The authors declare no competing interests. or product names is for descriptive purposes only and does not imply endorsement by the US Government. Additional information Supplementary information is available for this paper at https://doi.org/10.1038/ Author contributions s41561-018-0277-3. S.L.L., A.A.B., E.G.P. and J.R.V. carried out the organic isotope analyses, S.L.L. and Reprints and permissions information is available at www.nature.com/reprints. A.A.B. carried out the biomarker analyses, S.L.L. interpreted the biomarker and Correspondence and requests for materials should be addressed to S.L.L. 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 Publisher’s note: Springer Nature remains neutral with regard to jurisdictional claims in mixing model, K.H.F., J.C.Z., S.M.T., L.R.K., T.J.B. and A.A.B. contributed to major published maps and institutional affiliations. improvements within the models and data interpretation, S.L.L. wrote the paper, and all © The Author(s), under exclusive licence to Springer Nature Limited 2018

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Methods chloride (aromatic fraction) and 70% methylene chloride:30% MeOH (v:v) Organic δ13C measurements. Sediment samples were powdered via a ball mill, and (polar fraction). Aliphatic fractions were analysed via gas chromatography–mass 1.5 g of powdered sediment were decarbonated using 15 ml of 1 N HCl for 30 min spectrometry on a Thermo Scientific Trace 1310 Gas Chromatograph coupled to to 2 h to remove carbonate. Samples were neutralized via repeated rinsing with a thermo Scientific ISQ LT single quadrupole mass spectrometer. A Restek Rxi- 13 5HT fused silica column (30 m length ​ 0.25 mm internal diameter ​ 0.25 ​m film MilliQ water until the pH rose above 5. %TOC and bulk δ​ Corg were measured × × µ using a Costech ECS 4010 Elemental Analyzer coupled to a Termo Delta Plus thickness) was used with a helium carrier gas. The oven programme started with –1 isotope ratio mass spectrometer. Te combustion reactor (length 45.4 cm, diameter an injection temperature of 40 °C held for 1.5 min, followed by a rise of 15.0 °C min 18 mm) was packed with chromium(iii) oxide and silvered cobalt(ii,iii) and until the temperature reached 140 °C, after which the temperature increased at –1 heated to 650 °C, and the oxidation reactor (length 45.4 cm, diameter 14 mm) was 6.0 °C min until it reached 320 °C, at which it was held for 20 min, for a total run packed with reduced copper wires and heated to 1,020 °C. Te water trap (length time of 58 min. The flow was split and 90% was transferred to an ISQ quadrupole

11 cm, diameter 8 mm) was packed with magnesium perchlorate. CO2 and N2 were mass spectrometer with a transfer line temperature of 340 °C and electron separated on a gas chromatography column heated to 50 °C prior to transfer to ionization of 300 °C, which was scanned over the mass range 50–550 amu at –1 the isotope ratio mass spectrometer. Samples were weighed into tin capsules and 5 scans s . The other 10% of the split flow was routed to a flame ionization detector analysed with a suite of blanks and fve laboratory standards with known isotope held at 330 °C. Biomarker quantifications were determined from flame ionization 13 detector peak area, using AGSO standard oil, Macondo oil and an n-C to and %TOC values (Supplementary Table 4). δ​ Corg values and %TOC of the 10 13 n-C alkane standard and response curve. Mass spectra were used for compound samples were calculated using a three-point standard curve, and δ​ Corg values are 40 reported in standard delta notation against Vienna Pee Dee Belemnite. Instrument identification by comparison with standards and published spectra. precision and accuracy were ±​0.18‰ and ±​0.61‰, respectively. Duplicate samples 13 Code availability. The mixing model computer code is available from the authors had a reproducibility of ±​0.2‰ for δ​ Corg and ±​0.05% for %TOC. upon reasonable request. 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 Data availability of N2. Lipid extracts were separated into fractions (aliphatic, aromatic and polar) The authors declare that the data supporting the findings of this study are available using gravity column separations with a fine mesh silica stationary phase and within the paper and its Supplementary Information files. Additional source data mobile phases of 100% hexane (aliphatic fraction), 90% hexane 10% methylene for Fig. 4 is available upon reasonable request.

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