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ABSTRACT Records suggest that the Permo–Triassic mass extinction (PTME) involved one of the most severe terrestrial ecosystem collapses of the Phanerozoic. However, it has proved difficult

to constrain the extent of the primary productivity loss on land, hindering our understanding of the effects on global biogeochemistry. We build a new biogeochemical model that couples the

global Hg and C cycles to evaluate the distinct terrestrial contribution to atmosphere–ocean biogeochemistry separated from coeval volcanic fluxes. We show that the large short-lived Hg

spike, and nadirs in δ202Hg and δ13C values at the marine PTME are best explained by a sudden, massive pulse of terrestrial biomass oxidation, while volcanism remains an adequate explanation

for the longer-term geochemical changes. Our modelling shows that a massive collapse of terrestrial ecosystems linked to volcanism-driven environmental change triggered significant

biogeochemical changes, and cascaded organic matter, nutrients, Hg and other organically-bound species into the marine system. SIMILAR CONTENT BEING VIEWED BY OTHERS ASSESSING THE IMPORTANCE

OF THERMOGENIC DEGASSING FROM THE KAROO LARGE IGNEOUS PROVINCE (LIP) IN DRIVING TOARCIAN CARBON CYCLE PERTURBATIONS Article Open access 28 October 2021 SIX-FOLD INCREASE OF ATMOSPHERIC

_P_CO2 DURING THE PERMIAN–TRIASSIC MASS EXTINCTION Article Open access 09 April 2021 PALEOCENE/EOCENE CARBON FEEDBACKS TRIGGERED BY VOLCANIC ACTIVITY Article Open access 31 August 2021

INTRODUCTION The Permo–Triassic mass extinction (PTME) is the largest known extinction in Earth′s history, with the loss of ~90% of species in the sea and ~70% of species on land1,2,3,4. The

PTME has been causally linked to the emplacement of the Siberian Traps Large Igneous Province (LIP) and associated volcanic gas emissions (especially CO2, SO2 and halogens), via widespread

environmental changes such as warming and oceanic anoxia5,6,7,8. The PTME also saw a crisis in terrestrial ecosystems, with loss of plant diversity, increased wildfire activity and

consequent enhanced soil erosion9,10,11,12,13,14. Recent work has shown that the disruption of vegetation started before and culminated at the marine extinction level (Fig. 1), implying that

the environmental disaster impacted terrestrial ecosystems first12,14. The cause of the terrestrial mass extinction is still unclear, and several kill mechanisms have been hypothesised. For

example, a shift from a humid warm climate to an unstable highly seasonal climate and an associated increase in wildfires affected the equatorial Permo–Triassic peatlands, drastically

reducing the abundance and diversity of the flora14; abnormal pollen and spores found in different localities around the world during the PTME interval suggest widespread mutagenesis

possibly linked to an increase in UV-B radiation due to ozone depletion15,16; a terrestrial S-isotope record from the Karoo basin in South Africa could indicate volcanically driven acid rain

at the P–T transition17 that might have also severely impacted the flora. Whilst the taxonomic losses in terrestrial ecosystems are becoming clearer12,18, and local, enhanced input of

terrestrial material into marine environments has been recorded9,11, the biogeochemical impacts and feedbacks on the exogenic C cycle are not known. It is possible that these impacts were

severe: the PTME represents the largest, and maybe the only, known mass extinction of insects19, suggesting that there may have been a substantial decrease in available food sources at the

lowest levels of the food chain. The PTME is marked by an approximately two- to four-fold increase in marine sedimentary Hg concentration with respect to background levels during a ~400 kyr

interval also characterised by negative δ13C values20, which implies a relatively long-term injection of Hg and 13C-depleted CO2 into the atmosphere–land–ocean system during this time.

Superimposed on this trend is a prominent, short-lived Hg spike, which is usually expressed as Hg/TOC, given the affinity of Hg with organic matter, and which is coincident with the collapse

of the terrestrial ecosystems14, the onset of the marine mass extinction interval and a sharp minimum in δ13C values (Fig. 1). While higher Hg/TOC values have been reported preceding the

marine extinction in the deep-water settings of Japan21, these are an artefact of normalisation to values of TOC which are below the analytical detection limit (<<0.1%), and do not

track increases in Hg concentrations (see TOC and Hg data in the Supplementary Information of ref. 21). The Hg record is interpreted as evidence of increased Hg input into the Earth's

surface system from the Siberian Traps13,21,22,23,24. However, Hg is also stored in large quantities in terrestrial biomass and soils25,26,27. The mobilisation of these terrestrial

reservoirs during the PTME could also result in the increased loading of Hg to aquatic environments, even without an elevated volcanic Hg flux to the atmosphere23,27,28,29. Any such changes

in the soil and biomass carbon reservoir would also have direct implications for the release of C to the atmosphere and the sedimentary δ13C record. Hg isotopes can be used to better

understand how Hg has been transported into the sedimentary environment. Mass-independent fractionation (MIF—denoted by Δ) occurs due to aqueous photoreduction of Hg2+ to

\({\mathrm{Hg}}_{({\mathrm{g}})}^0\) that takes place in surface waters and in clouds29. This results in positive Δ199Hg values in the remaining water Hg2+ pool, hence positive Δ199Hg values

in sediments dominated by atmospheric Hg2+ deposition. Conversely, slightly negative Δ199Hg values characterise the terrestrial reservoir (soil and biomass), which primarily captures

\({\mathrm{Hg}}_{({\mathrm{g}})}^0\)29. During Hg uptake by plants additional mass-dependent fractionation (MDF) and MIF occur, resulting in more negative δ202Hg and Δ199Hg values29. The

Permian–Triassic boundary shallow-water record of Meishan shows a prominent negative δ202Hg excursion in correspondence to the Hg spike coupled to a small negative Δ199Hg shift, but deeper

water successions show persistent positive Δ199Hg20,21,23,24. Therefore, isotope data appear to indicate that the Hg was transported to the deep-water settings mostly via the atmosphere, and

to the shallow-water settings via continental runoff and via the atmosphere20,21,23,24. The δ13C records at the Permian–Triassic boundary show two minima30,31, which are here called EP. I

and II (EP. = episode) following ref. 30 (Fig. 1). The observed δ13C trends are similarly recorded in different depositional settings and by different substrates (carbonates, bulk organic

matter, separate plant remains)14,30,31,32, strongly indicating that they represent actual changes in the C-isotope composition of the reservoirs of the exogenic C cycle. The Hg

concentration spike occurs in the same interval as the minimum in δ13C values associated with the PTME (EP. I, Fig. 1). At Meishan, where the chronostratigraphic framework is well

established33,34, the initial large Hg spike occurs about 60 Kyrs after the onset of the carbon-isotope perturbation (Fig. 1). Data from non-marine end-Permian successions confirms this

diachrony13,14 (Fig. 1) and show that the Hg spike is also coincident with a sudden decrease in Total Organic Carbon (TOC) values to almost zero14 (Fig. 1). The combination of geochemical

and palaeontological data from these sections shows that the terrestrial ecosystem disruption started with the onset of the carbon-isotope perturbation and climaxed at the very sharp δ13C

minimum (EP. I), coincident with the Hg spike (Fig. 1), and the start of the main marine extinction interval. This mismatch in both timings and fluxes between the C and Hg cycles at the

PTME, suggest that the 13C-depleted C and the Hg came from multiple sources. Overall, the interplay between volcanism and terrestrial reservoir changes in controlling PTME biogeochemistry is

not well known, and previous attempts to model the δ13C record have been fundamentally hampered because of the lack of an independent tracer of the C source. To overcome this, we use

published records of δ13C and Hg systematics to jointly constrain a new coupled C and Hg biogeochemical model. Our model shows that the large, sudden geochemical shifts at the PTME are best

explained by a massive pulse of terrestrial biomass oxidation, while Siberian Traps volcanism can explain the longer-term geochemical changes. RESULTS AND DISCUSSION A COUPLED HG-C CYCLE

MODEL Figure 2 shows the biogeochemical box model developed here. The full model derivation follows in the ‘Methods’ section. The model combines a multi-box sediment–ocean–atmosphere

carbon-alkalinity cycle (based on previous work35,36,37), with the global mercury cycle38,39. The ocean is split into ‘surface’, ‘high-latitude’ and ‘deep’ boxes. It considers ocean

circulation and carbonate speciation, and contains a simplified organic carbon cycle in which burial rates are prescribed. As well as computing the global C and Hg cycles it also computes

δ13C of all C reservoirs and δ202Hg of the ocean reservoirs. A full atmosphere–ocean model of δ202Hg would require dynamic biosphere reservoirs, which would greatly increase model

complexity. We therefore simplify the system to a mixing model for marine δ202Hg, in which atmospheric and riverine inputs have different isotopic signatures. The atmosphere is assumed to

have \(\delta ^{202}{\mathrm{Hg}}_{{\mathrm{atm}}} = - 1\) ‰, and riverine input is assumed to have \(\delta ^{202}{\mathrm{Hg}}_{{\mathrm{runoff}}} = - 3\) ‰, following ref. 40. The model

is set up for the late Permian by reducing the solar constant to that of 250 Ma, and increasing the background tectonic CO2 degassing rate to 1.5 times the present day (_D_ = 1.5), in line

with estimates for the Late Permian41. To obtain the observed pre-event ocean–atmosphere δ13C composition of ~3.5‰, we set the rate of land-derived organic carbon burial to 20% higher than

present day, and adjust the composition of the weathered carbonate reservoir to 3‰. This is consistent with high terrestrial productivity and C burial in the Permian (e.g., coal forests and

mires) and rapid recycling of more recently buried and 13C-enriched carbonate material. In the following paragraphs, we test two model end-member scenarios: (I) the release of volcanic and

thermogenic Hg and C from Siberian Traps activity alone, and (II) with the additional release of Hg and C as a consequence of the collapse of the terrestrial ecosystems. VOLCANIC AND

THERMOGENIC DEGASSING We first model the release of volcanic/volcanogenic Hg and C from the Siberian Traps. Existing radioisotope data show that the extinction, the negative δ13C excursion

and Hg spike might have all occurred during the intrusive phase of the Siberian Traps20,34,42. It is suggested that the emplacement of large sills caused the combustion or thermal

decomposition of organic-rich sediments with the consequent release of thermogenic volatiles, such as C and Hg20,22. It has been proposed that over a ~400 Kyrs intrusive phase the Siberian

Traps emitted ~7600–13,000 Mg yr−1 of volcanic Hg, which included both magmatic and coal-derived Hg20,22,23. Relating this Hg release to the background volcanic source is not straightforward

because estimates of the background source vary, but taking the most likely present day range25 (~90–360 Mg yr−1), and further constraining this by taking into account the need to balance

overall sedimentary burial of Hg (~190 Mg yr−1), and the overall ~50% increase in tectonic degassing in the late Palaeozoic relative to today43, we arrive at a best guess for the background

late Permian Hg flux of ~300 Mg yr−1. This means that the Siberian Traps eruption increased the geogenic Hg input by a factor of ~25–43 over ~400 Kyrs. To test this scenario, we model

Siberian Traps intrusion by increasing the volcanic Hg source by 25–43-fold for 450 Kyrs, while also increasing the CO2 source in line with estimates44,45 for Siberian Traps degassing based

on magma volumes and sediment intrusion (by 4-8 × 1012 moles/year). The CO2 released by contact metamorphism at the PTME is assumed to have an average δ13C composition of −25‰44.

Specifically, the input functions are: $$f_{CO_{2input}} = \, \left[ { - 253 - 251.99 - 251.98 - 251.56 - 251.55 - 251} \right],\\ \big[ {0\,0\,CO_{2ramp}\,CO_{2ramp}\,0\,0} \big]$$

$$F_{Hg_{input}} = \, \left[ { - 253 - 251.99 - 251.98 - 251.56 - 251.55 - 251} \right],\\ \big[ {1\,1\,Hg_{ramp}\,Hg_{ramp}\,1\,1} \big]$$ Here the first vector is time in millions of years

and the second is the flux alteration at that time. Here \(CO_{2ramp}\) is the additional CO2 release in mol yr−1, and \(Hg_{ramp}\) is the relative Hg degassing rate increase. For the

duration of these pulses, the thermohaline circulation is also assumed to collapse due to warming and freshwater input46. We reduce the circulation term to 1 Sv over this period, which

allows more rapid change in the model surface ocean C isotopes and Hg loading. This is a large reduction, and also reflects the simple structure of the model in which the entire low-latitude

surface ocean is represented by a single box, and so is well-mixed. Figure 3a–d shows that this magnitude and timing of release of C and Hg is capable of driving the longer-term decline in

carbonate δ13C, and the coeval long-term approximately two- to four-fold enrichment in shallow sediment Hg/TOC that is observed in Meishan. However, the model scenario does not capture the

spike in Hg concentration, or nadir in δ13C (EP. I30 in Fig. 1) that are coincident with the final stage of the terrestrial extinction. It also does not produce any substantial change in

marine δ202Hg isotopes (Fig. 1d), because the primary Hg source to the ocean is the atmosphere for the full model run. Within the model, we have also explored a scenario wherein the large Hg

pulse represents a further rapid pulse of LIP volcanism. We have attempted this scenario in Supplementary Note 1 (Scenario I–2), where a 1 Kyr volcanic pulse is assumed to raise the Hg and

C input rates by a further factor of 5. While the Hg/TOC can indeed be explained by an additional short-lived pulse of Hg, we require the total release rate of Hg to be ~200 times greater

than background levels, and even then, this scenario fails to reproduce any of the Hg isotope signature or the nadir in carbonate δ13C (Supplementary Fig. 1). TERRESTRIAL ECOSYSTEM COLLAPSE

For scenario II, we explore the additional effects of a geologically rapid (~1 Kyr) pulse of Hg and C as the result of the collapse of terrestrial ecosystems at the PTME. The magnitude of

this Hg flux is again difficult to quantify precisely, and we explore an increase of 100-fold over background conditions. This level of increase represents the magnitude required to drive

the sedimentary Hg signal that we observe, and is compatible with the available terrestrial biosphere Hg reservoir: total soil Hg is estimated to be on the order of ~106 Mg Hg when

considering a soil depth of ~15 cm47. So, our model Hg delivery flux would require decimetre-scale soil organic matter oxidation over 1000 years, coincident with the PTME and the sharp EP. I

negative δ13C shift11. The Hg pulse is delivered directly to the low-latitude surface ocean via runoff in the model, and is accompanied by a pulse of ‘soil oxidation’ C which we assume

raises the global rate of oxidative weathering by a factor of 30—a number chosen to have the observed level of impact on the C-isotope record while being compatible with the Hg input change.

We also assume a cessation of terrestrial organic C burial. Terrestrial Hg deposition and erosion is not altered during the pulse as the fluxes are minor by comparison. The new model

functions applied in addition to the longer-term inputs of scenario I are: $$F_{Cburial} = \, \left[ { - 253 - 251.951 - 251.950 - 251.949 - 251.948 - 251} \right],\\ \big[

{1\,1\,C_{ramp}\,C_{ramp}\,1\,1} \big]$$ $$F_{oxidw} = \, \left[ { - 253 - 251.951 - 251.950 - 251.949 - 251.948 - 251} \right],\\ \big[ {1\,1\,O_{ramp}\,O_{ramp}\,1\,1} \big]$$ $$F_{runoff}

= \, \left[ { - 253 - 251.951 - 251.950 - 251.949 - 251.948 - 251} \right],\\ \big[ {1\,1\,Hg_{bio}\,Hg_{bio}\,1\,1} \big]$$ Here, the first vector is time in millions of years, and the

second shows flux multipliers at these times. \(C_{ramp},O_{ramp}\) and \(Hg_{bio}\) denote the relative rate of land organic C burial, oxidative weathering and Hg runoff, respectively, and

are set at 0, 30 and 100, respectively, for the duration of the 1-kyr pulse. This ‘biosphere’ pulse causes a short-term large concentration spike in the shallow marine Hg reservoir and its

sediments, which is superimposed on the volcanically driven changes (Fig. 3e–h). The Hg spike is far larger than would be expected from simply increasing the volcanic source by the same

amount because the biospheric Hg is delivered directly to the surface ocean and sedimentation occurs mostly on the shelf. With the inclusion of terrestrial C oxidation and cessation of

terrestrial carbon burial, the model also replicates the transient shift to more negative δ13C values recorded at the marine extinction interval (EP. I30,31 in Fig. 1): Terrestrial C

oxidation is a source of isotopically light C48. The model now also shows a sharp negative δ202Hg shift in the shallow ocean box, which is triggered by increased Hg riverine input, but shows

no change in the deeper ocean, where the source of Hg remains predominantly atmospheric. This also compares well with existing records (Fig. 3). At Meishan, which was located in the margins

of the Yangtze carbonate platform, the Hg and Hg/TOC spike is coincident with more negative δ202Hg values (Fig. 1), while in the deeper water sections of south China the values are more

positive20,23. Hence, oxidation of terrestrial biomass is a compelling scenario to explain the palaeontological, sedimentological and geochemical data. There is clear observational evidence

for the collapse of the terrestrial ecosystems and cessation of terrestrial C burial, stratigraphic evidence supporting the sequence and timing of the events (onset of the δ13C

shift—collapse of the terrestrial ecosystem—Hg and C spike), sufficient quantity of Hg available, consistency with the isotopic evidence for changing Hg sources, and consistency with the

δ13C records. MASSIVE TERRESTRIAL BIOMASS OXIDATION DURING THE PTME Using our coupled C–Hg biogeochemical model, we show that the massive collapse of terrestrial ecosystems and oxidation of

terrestrial biomass during the Permian–Triassic extinction had a huge impact on global Hg and C biogeochemistry. Hg stored in the terrestrial reservoirs was rapidly released as a consequence

of the loss of terrestrial biomass and increased soil erosion9,14. This mechanism is the best explanation for the sharp increased loading of Hg into both terrestrial and marine water bodies

and the negative shift in δ202Hg in coincidence with the marine mass extinction. Contemporaneously, increased soil carbon oxidation introduced large quantities of isotopically light C,

accounting for the sharp negative δ13C anomaly registered in the sedimentary record (EP. I30). In the model, the emission of Hg and C from magma and heating of sedimentary organic matter

during the intrusive phase of the Siberian Traps LIP emplacement can account for the smaller, two- to fourfold increase of Hg concentrations with respect to background levels, and the

relatively longer negative δ13C trend that is recorded by both carbonates and organic matter, in marine and terrestrial settings. A new scenario emerges for the PTME that links the collapse

of ecosystems on land to the global geochemical changes recorded at the marine extinction interval. The disruption of terrestrial environments started during the initial phases of the

Siberian Traps emplacement likely due to the release of volcanic gases as CO2, SO2 and halogens, which could have triggered acid rain, ozone depletion, volcanic darkness, rapid cooling and

subsequent global warming8,49. At the culmination of the terrestrial disturbance interval, when the ecosystems totally collapsed, large amounts of 13C-depleted C and Hg deriving from a

massive oxidation of terrestrial biomass were transported into aqueous habitats causing a steep decline in sedimentary δ13C (carbonates and organic matter), a sedimentary Hg concentrations

spike and a shift in δ202Hg (Fig. 3). At this level, the marine mass extinction started. This, according to the existing chronostratigraphic framework, happened ∼60 Kyrs after the onset of

the carbon-isotope perturbation and of the terrestrial ecological disturbances14 (Fig. 1). The biogeochemical cycle of Hg is intimately linked to the cycle of organic matter and its

constituting elements, such as C, N, S and P50. Hence, besides Hg and C, other organically-bound species would have been transferred from the terrestrial reservoirs into the marine system in

large quantities at EP. I (Fig. 1). Addition of these species, particularly the nutrients P and N, are easily capable of driving ecosystem turnover, anoxia and eutrophication, and it is

likely that this terrestrial input contributed to the marine extinction9,11. Our model does not include these additional cycles, but other models have shown that a relatively small increase

in marine P delivery (2–3-fold) has the potential to drive marine anoxia or euxinia51,52. The scale of the terrestrial ecosystem collapse at the PTME could explain the severity of the biotic

crisis at the Permian–Triassic boundary at all trophic levels, and should be a key consideration for future research. For other events, the Hg records are not so consistent nor as detailed

as for the PTME. However, it is very likely that future research on other intervals could show the same Hg and C patterns as for the PTME. METHODS MODEL DERIVATION This model is designed to

track the transfer and isotopic signature of atmospheric and marine carbon and mercury over geological time, while being broadly applicable to changes on the timescale of ocean circulation.

The biogeochemical system is taken largely from ref. 36, with some additions from refs. 37,53,54, with the underlying hydrological model from ref. 35. The Hg cycle follows ref. 25. MODEL

STRUCTURE The model has three ocean boxes: surface (s), high latitude (h) and deep (d). As in ref. 35, the surface box is 100-m deep and occupies 85% of the ocean surface, whereas the

high-latitude box is 250-m deep and represents 15% of the ocean surface. Each ocean box includes the same biogeochemical species, and a thermohaline circulation mixes the boxes in the order

s, h, d. The upper boxes exchange with the atmosphere, which is a single box. As well as transfer fluxes between ocean and atmosphere boxes, biogeochemical fluxes of weathering, degassing

and burial operate between the surface system and crust. MODEL SPECIES All model species are shown in Table 1. MODEL FLUXES Model fluxes, with equations and present values are shown in Table

 2. NON-FLUX CALCULATIONS Atmospheric CO2 volume ratio is calculated as: $${\mathrm{CO}}_2{\mathrm{ppm}} = 280\frac{{{\mathrm{CO}}_{2{\mathrm{a}}}}}{{{\mathrm{CO}}_{2{\mathrm{a}}_0}}}$$

where \({\mathrm{CO}}_{2{\mathrm{a}}}\) is atmospheric CO2 in moles, and \({\mathrm{CO}}_{2{\mathrm{a}}_0}\) is this value at present day. Global average surface temperature (GAST) is:

$${\mathrm{GAST}} = 288 + {\mathrm{k}}_{{\mathrm{clim}}}\left( {\frac{{{\mathrm{log}}\left( {\frac{{{\mathrm{CO}}_2{\mathrm{ppm}}}}{{280}}} \right)}}{{\log \left( 2 \right)}}} \right) -

7.4\left( {\frac{{{\mathrm{t}}_{{\mathrm{geol}}}}}{{ - 570}}} \right)$$ where kclim is climate sensitivity to doubling CO2, and tgeol is time in millions of years before present and is

expressed in negative terms. Low-latitude surface temperature (Ts) is assumed to scale by \({\textstyle{2 \over 3}}\) times global temperature change, and both high-latitude (Th) and deep

(Td) temperature are assumed to follow global temperature change. For carbonate speciation, effective equilibrium constants are calculated following refs. 36,55:

$${\mathrm{K}}_{{\mathrm{carb}}} = 5.75 \times 10^{ - 4} + 6 \times 10^{ - 6}({\mathrm{T}}_{\mathrm{j}} - 278)$$ $${\mathrm{K}}_{{\mathrm{CO}}_2} = 0.035 + 0.0019({\mathrm{T}}_{\mathrm{j}} -

278)$$ Dissolved carbon species are then calculated following Walker and Kasting36: $$\left[ {{\mathrm{HCO}}_3^ - } \right]_{\mathrm{j}} = \frac{{{\mathrm{DIC}}_{\mathrm{j}} - \sqrt

{{\mathrm{DIC}}_{\mathrm{j}}^2 - {\mathrm{ALK}}_{\mathrm{j}}\left( {2{\mathrm{DIC}}_{\mathrm{j}} - {\mathrm{ALK}}_{\mathrm{j}}} \right)\left( {1 - 4{\mathrm{K}}_{{\mathrm{carb}}}} \right)}

}}{{1 - 4{\mathrm{K}}_{{\mathrm{carb}}}}}$$ $$\left[ {{\mathrm{CO}}_3^{2 - }} \right]_{\mathrm{j}} = \frac{{{\mathrm{ALK}}_{\mathrm{j}} - \left[ {{\mathrm{HCO}}_3^ - }

\right]_{\mathrm{j}}}}{2}$$ $${\mathrm{pCO}}_{2{\mathrm{j}}} = \frac{{{\mathrm{K}}_{{\mathrm{CO}}_{2}}\left[ {{\mathrm{HCO}}_3^ - } \right]^2}}{{[{\mathrm{CO}}_3^{2 - }]}}$$ We also

explicitly calculate [H+] concentration to observe model pH: $$\left[ {{\mathrm{H}}^ + } \right] = {\mathrm{K}}_2\frac{{\left[ {{\mathrm{HCO}}_3^ - } \right]}}{{[{\mathrm{CO}}_3^{2 - }]}}$$

Calcium carbonate saturation state is calculated as: $${\mathrm{\Omega }}_{\mathrm{j}} = \frac{{\left[ {{\mathrm{Ca}}} \right]_{\mathrm{j}}\left[ {{\mathrm{CO}}_3^{2 - }}

\right]_{\mathrm{j}}}}{{{\mathrm{K}}_{{\mathrm{sp}}}}}$$ where Ωj is the CaCO3 saturation state in box j, and Ksp is the solubility product. For terrestrial chemical weathering, temperature

dependence of basalt and granite weathering is calculated as: $${\mathrm{f}}_{{\mathrm{Tbas}}} = {\mathrm{e}}^{0.0608\left( {{\mathrm{GAST}} - 288} \right)}\left( {1 + 0.038\left(

{{\mathrm{GAST}} - 288} \right)} \right)^{0.65}$$ $${\mathrm{f}}_{{\mathrm{Tgran}}} = {\mathrm{e}}^{0.0724\left( {{\mathrm{GAST}} - 288} \right)}\left( {1 + 0.038\left( {{\mathrm{GAST}} -

288} \right)} \right)^{0.65}$$ And temperature dependence of carbonate weathering: $${\mathrm{f}}_{{\mathrm{Tcarb}}} = 1 + 0.087({\mathrm{GAST}} - 288)$$ FIXED PARAMETERS Fixed parameters

are shown in Table 3. DIFFERENTIAL EQUATIONS The following equations track the 11 non-water species from Table 1. Atmospheric CO2: $$\frac{{\mathrm{d}}\left( {{\mathrm{CO}}_{2{\mathrm{a}}}}

\right)}{{{\mathrm{dt}}}} = - {\mathrm{f}}_{{\mathrm{airsea}}_{\mathrm{s}}} - {\mathrm{f}}_{{\mathrm{airsea}}_{\mathrm{h}}} + {\mathrm{f}}_{{\mathrm{ccdeg}}} +

{\mathrm{f}}_{{\mathrm{ocdeg}}} + {\mathrm{f}}_{{\mathrm{oxidw}}} - {\mathrm{f}}_{{\mathrm{locb}}} - {\mathrm{f}}_{{\mathrm{carbw}}} - 2{\mathrm{f}}_{{\mathrm{silw}}} +

{\mathrm{f}}_{{\mathrm{CO}}_{2\mathrm{input}}}$$ Low-latitude surface ocean DIC: $$\frac{{\mathrm{d}}\left( {{\mathrm{DIC}}_{\mathrm{s}}} \right)}{{{\mathrm{dt}}}} =

{\mathrm{f}}_{{\mathrm{airsea}}_{\mathrm{s}}} + {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{ds}}} - {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{sh}}} + 2{\mathrm{f}}_{{\mathrm{carbw}}} +

2{\mathrm{f}}_{{\mathrm{silw}}} - {\mathrm{f}}_{{\mathrm{mccb}}} - {\mathrm{f}}_{{\mathrm{mocb}}}$$ High-latitude surface ocean DIC: $$\frac{{\mathrm{d}}\left( {{\mathrm{DIC}}_{\mathrm{h}}}

\right)}{{{\mathrm{dt}}}} = {\mathrm{f}}_{{\mathrm{airsea}}_{\mathrm{h}}} + {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{sh}}} - {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{hd}}}$$ Deep ocean DIC:

$$\frac{{\mathrm{d}}\left( {{\mathrm{DIC}}_{\mathrm{d}}} \right)}{{{\mathrm{dt}}}} = {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{hd}}} - {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{ds}}}$$

Low-latitude surface ocean alkalinity: $$\frac{{{\mathrm{d}}\left( {{\mathrm{ALK}}_{\mathrm{s}}} \right)}}{{{\mathrm{dt}}}} = {\mathrm{tran}}_{{\mathrm{ALK}}_{\mathrm{ds}}} -

{\mathrm{tran}}_{{\mathrm{ALK}}_{\mathrm{sh}}} + 2{\mathrm{f}}_{{\mathrm{carbw}}} + 2{\mathrm{f}}_{{\mathrm{silw}}} - 2{\mathrm{f}}_{{\mathrm{mccb}}}$$ High-latitude surface ocean

alkalinity: $$\frac{{\mathrm{d}}\left( {{\mathrm{ALK}}_{\mathrm{h}}} \right)}{{{\mathrm{dt}}}} = {\mathrm{tran}}_{{\mathrm{ALK}}_{\mathrm{sh}}} -

{\mathrm{tran}}_{{\mathrm{ALK}}_{\mathrm{hd}}}$$ Deep ocean alkalinity: $$\frac{{\mathrm{d}}\left( {{\mathrm{ALK}}_{\mathrm{d}}} \right)}{{{\mathrm{dt}}}} =

{\mathrm{tran}}_{{\mathrm{ALK}}_{\mathrm{hd}}} - {\mathrm{tran}}_{{\mathrm{ALK}}_{\mathrm{ds}}}$$ δ13C of atmospheric CO2: $$\frac{{\mathrm{d}}\left( {\delta

^{13}{\mathrm{CO}}_{2{\mathrm{a}}} \cdot {\mathrm{CO}}_{2{\mathrm{a}}}} \right)}{{{\mathrm{dt}}}} = - {\mathrm{f}}_{{\mathrm{airsea}}_{\mathrm{s}}}\delta ^{13}{\mathrm{C}}_{{\mathrm{atm}}} -

{\mathrm{f}}_{{\mathrm{airsea}}_{\mathrm{h}}}\delta ^{13}{\mathrm{C}}_{{\mathrm{atm}}} + {\mathrm{f}}_{{\mathrm{ccdeg}}}\delta ^{13}{\mathrm{C}}_{\mathrm{C}} \\ +

{\mathrm{f}}_{{\mathrm{ocdeg}}}\delta ^{13}{\mathrm{C}}_{\mathrm{G}} + {\mathrm{f}}_{{\mathrm{oxidw}}}\delta ^{13}{\mathrm{C}}_{\mathrm{G}} - {\mathrm{f}}_{{\mathrm{locb}}} \left( {\delta

^{13}{\mathrm{C}}_{{\mathrm{atm}}}\, - \Delta {\mathrm{B}}} \right) \\ - {\mathrm{f}}_{{\mathrm{carbw}}}\delta ^{13}{\mathrm{C}}_{{\mathrm{atm}}} - 2{\mathrm{f}}_{{\mathrm{silw}}}\delta

^{13}{\mathrm{C}}_{{\mathrm{atm}}} + {\mathrm{f}}_{{\mathrm{CO}}_{2{\mathrm{input}}}}\delta ^{13}{\mathrm{C}}_{{\mathrm{input}}}$$ δ13C of low-latitude surface ocean DIC:

$$\frac{{\mathrm{d}}\left( {\delta} ^{13}{\mathrm{DIC}}_{\mathrm{s}} \cdot {\mathrm{DIC}}_{\mathrm{s}} \right)}{{{\mathrm{dt}}}} = \, {\mathrm{f}}_{{\mathrm{airsea}}_{\mathrm{s}}}{\delta}

^{13}{\mathrm{C}}_{{\mathrm{atm}}} + {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{ds}}}{\delta} ^{13}{\mathrm{DIC}}_{\mathrm{d}} \, - {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{sh}}}{\delta}

^{13}{\mathrm{DIC}}_{\mathrm{s}} \\ + {\mathrm{f}}_{{\mathrm{carbw}}}{\delta} ^{13}{\mathrm{C}}_{{\mathrm{atm}}} + {\mathrm{f}}_{{\mathrm{carbw}}}{\delta} ^{13}{\mathrm{C}}_{\mathrm{C}}\, +

2{\mathrm{f}}_{{\mathrm{silw}}}{\delta} ^{13}{\mathrm{C}}_{{\mathrm{atm}}} \\ - {\mathrm{f}}_{{\mathrm{mccb}}}{\delta} ^{13}{\mathrm{DIC}}_{\mathrm{s}} -

{\mathrm{f}}_{{\mathrm{mocb}}}({\delta} ^{13}{\mathrm{DIC}}_{\mathrm{s}} - {\Delta} {\mathrm{B}})$$ δ13C of high-latitude surface ocean DIC: $$\frac{{{\mathrm{d}}\left( {{\updelta

}}^{13}{\mathrm{DIC}}_{\mathrm{h}} \cdot {\mathrm{DIC}}_{\mathrm{h}} \right)}}{{{\mathrm{dt}}}} = {\mathrm{f}}_{{\mathrm{airsea}}_{\mathrm{h}}}{\updelta}^{13}{\mathrm{C}}_{{\mathrm{atm}}} +

{\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{sh}}}{\updelta}^{13}{\mathrm{DIC}}_{\mathrm{s}} - {\mathrm{tran}}_{{\mathrm{DIC}}_{\mathrm{hd}}}{\updelta}^{13}{\mathrm{DIC}}_{\mathrm{h}}$$ δ13C of

deep ocean DIC: $$\frac{{\mathrm{d}}\left( {{\updelta }}^{13}{\mathrm{DIC}}_{\mathrm{d}} \cdot {\mathrm{DIC}}_{\mathrm{d}} \right)}{{{\mathrm{dt}}}} =

{\mathrm{tran}}_{{\mathrm{DIC}}_{hd}}{\updelta}^{13}{\mathrm{DIC}}_{\mathrm{h}} - {\mathrm{tran}}_{{\mathrm{DIC}}_{ds}}{\updelta}^{13}{\mathrm{DIC}}_{\mathrm{d}}$$ Atmospheric Hg:

$$\frac{{{\mathrm{d}}\left( {{\mathrm{Hg}}_{\mathrm{a}}} \right)}}{{{\mathrm{dt}}}} =\, {\mathrm{f}}_{{\mathrm{Hg}}_{\mathrm{volc}}} + {\mathrm{f}}_{{\mathrm{Hg}}_{\mathrm{wildfire}}} -

{\mathrm{f}}_{{\mathrm{oceandep}}_{\mathrm{h}}} + {\mathrm{f}}_{{\mathrm{oceaneva}}_{\mathrm{h}}}\\ \, - \,{\mathrm{f}}_{{\mathrm{oceandep}}_{\mathrm{s}}} +

{\mathrm{f}}_{{\mathrm{oceaneva}}_{\mathrm{s}}} - {\mathrm{f}}_{{\mathrm{vegdep}}} + {\mathrm{f}}_{{\mathrm{vegeva}}}$$ Low-latitude surface ocean Hg: $$\frac{{{\mathrm{d}}\left(

{{\mathrm{Hg}}_{\mathrm{s}}} \right)}}{{{\mathrm{dt}}}} = {\mathrm{f}}_{{\mathrm{Hg}}_{\mathrm{runoff}}} + {\mathrm{f}}_{{\mathrm{oceandep}}_{\mathrm{s}}} -

{\mathrm{f}}_{{\mathrm{oceaneva}}_{\mathrm{s}}} + {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{ds}}} - {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{sh}}} - {\mathrm{f}}_{{\mathrm{Hg}}_{\mathrm{b}}}$$

High-latitude surface ocean Hg: $$\frac{{{\mathrm{d}}\left( {{\mathrm{Hg}}_{\mathrm{h}}} \right)}}{{{\mathrm{dt}}}} = {\mathrm{f}}_{{\mathrm{oceandep}}_{\mathrm{h}}} -

{\mathrm{f}}_{{\mathrm{oceaneva}}_{\mathrm{h}}} + {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{sh}}} - {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{hd}}}$$ Deep ocean Hg: $$\frac{{{\mathrm{d}}\left(

{{\mathrm{Hg}}_{\mathrm{d}}} \right)}}{{{\mathrm{dt}}}} = {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{hd}}} - {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{ds}}}$$ δ202Hg of Low-latitude surface ocean

Hg: $$\frac{{{\mathrm{d}}\left( {{\updelta }}^{202}{\mathrm{Hg}}_{\mathrm{s}} \cdot {\mathrm{Hg}}_{\mathrm{s}} \right)}}{{{\mathrm{dt}}}} = \,

{\mathrm{f}}_{{\mathrm{Hg}}_{\mathrm{runoff}}}{\updelta}^{202}{\mathrm{Hg}}_{{\mathrm{runoff}}} +

{\mathrm{f}}_{{\mathrm{oceandep}}_{\mathrm{s}}}{\updelta}^{202}{\mathrm{Hg}}_{{\mathrm{atm}}} - {\mathrm{f}}_{{\mathrm{oceaneva}}_{\mathrm{s}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{s}} \\ +

{\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{ds}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{d}} - {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{sh}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{s}} -

{\mathrm{f}}_{{\mathrm{Hg}}_{\mathrm{b}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{s}}$$ δ202Hg of High-latitude surface ocean Hg: $$\frac{{{\mathrm{d}}\left( {{\updelta

}}^{202}{\mathrm{Hg}}_{\mathrm{h}} \cdot {\mathrm{Hg}}_{\mathrm{h}} \right)}}{{{\mathrm{dt}}}} =

\,{\mathrm{f}}_{{\mathrm{oceandep}}_{\mathrm{h}}}{\updelta}^{202}{\mathrm{Hg}}_{{\mathrm{atm}}} - {\mathrm{f}}_{{\mathrm{oceaneva}}_{\mathrm{h}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{h}}\\

\, + {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{sh}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{s}} - {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{hd}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{h}}$$

δ202Hg of Deep ocean Hg: $$\frac{{{\mathrm{d}}\left( {{\updelta }}^{202}{\mathrm{Hg}}_{\mathrm{d}} \cdot {\mathrm{Hg}}_{\mathrm{d}} \right)}}{{{\mathrm{dt}}}} =

{\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{hd}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{h}} - {\mathrm{tran}}_{{\mathrm{Hg}}_{\mathrm{ds}}}{\updelta}^{202}{\mathrm{Hg}}_{\mathrm{d}}$$ DATA

AVAILABILITY The geochemical data used in this paper come from already published literature, as cited in the text. CODE AVAILABILITY MATLAB code to run the model is available from B.J.W.

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1 equilibrium. _Elsevier Oceanogr. Ser._ 65, 1–84 (2001). Article  Google Scholar  Download references ACKNOWLEDGEMENTS J.D.C. thanks Timothy M. Lenton for useful comments from which this

study emerged. J.D.C., R.J.N. and P.W. acknowledge support from NERC grant NE/P013724/1. J.D.C. also acknowledges the One Hundred Talent Program of China University of Geosciences (CUG)

Wuhan, China. B.J.W.M. acknowledges support from NERC grants NE/S009663/1 and NE/R010129/1 and from a University of Leeds Academic Fellowship. D.C., J.T., W.S. and Y.W. acknowledge National

Natural Science Foundation of China grants (grants 41530104, 41661134047). T.A.M. acknowledges funding from ERC consolidator grant (ERC-2018-COG- 818717 -V-ECHO). AUTHOR INFORMATION Author

notes * These authors contributed equally: Jacopo Dal Corso, Benjamin J. W. Mills. AUTHORS AND AFFILIATIONS * School of Earth and Environments, University of Leeds, Leeds, LS2 9JT, UK Jacopo

Dal Corso, Benjamin J. W. Mills, Robert J. Newton & Paul B. Wignall * State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences, Wuhan, 430074, China

Jacopo Dal Corso, Daoliang Chu, Wenchao Shu, Yuyang Wu & Jinnan Tong * Department of Earth Sciences, University of Oxford, South Parks Road, Oxford, OX1 3AN, UK Tamsin A. Mather Authors

* Jacopo Dal Corso View author publications You can also search for this author inPubMed Google Scholar * Benjamin J. W. Mills View author publications You can also search for this author

inPubMed Google Scholar * Daoliang Chu View author publications You can also search for this author inPubMed Google Scholar * Robert J. Newton View author publications You can also search

for this author inPubMed Google Scholar * Tamsin A. Mather View author publications You can also search for this author inPubMed Google Scholar * Wenchao Shu View author publications You can

also search for this author inPubMed Google Scholar * Yuyang Wu View author publications You can also search for this author inPubMed Google Scholar * Jinnan Tong View author publications

You can also search for this author inPubMed Google Scholar * Paul B. Wignall View author publications You can also search for this author inPubMed Google Scholar CONTRIBUTIONS J.D.C.

conceived the study. B.J.W.M. built the box model. J.D.C. and B.J.W.M. designed the model scenarios with in-depth inputs from T.A.M., P.B.W., D.C. and R.J.N. J.D.C., D.C., W.S., Y.W. and

P.B.W. compiled and discussed the geochemical and chronostratigraphic data. All authors discussed the results and contributed to the writing of the paper. P.B.W., R.J.N. and J.T. provided

the funding. CORRESPONDING AUTHORS Correspondence to Jacopo Dal Corso or Benjamin J. W. Mills. ETHICS DECLARATIONS COMPETING INTERESTS The authors declare no competing interests ADDITIONAL

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CITE THIS ARTICLE Dal Corso, J., Mills, B.J.W., Chu, D. _et al._ Permo–Triassic boundary carbon and mercury cycling linked to terrestrial ecosystem collapse. _Nat Commun_ 11, 2962 (2020).

https://doi.org/10.1038/s41467-020-16725-4 Download citation * Received: 23 December 2019 * Accepted: 18 May 2020 * Published: 11 June 2020 * DOI: https://doi.org/10.1038/s41467-020-16725-4

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