Atmospheric CO_2 drawdown during the Emeishan flood basalt volcanism

Domain: chemistry
Published: 2026-02-26
Editorial outcome: minor_revision

Abstract
The conventional model linking large igneous provinces (LIPs) to atmospheric CO concentrations (pCO ) predicts elevated CO during volcanic emplacement, leading to biotic crises. However, this generalization requires testing. Here, we reconstruct pCO variation throughout the Emeishan LIP emplacement (~260 Ma) using carbon isotopes from chlorophyll-derived compounds. Our high-resolution record reveals that pCO declined from ~700 ppm to ~350 ppm during the early and main flood basalt phases, then increased during subsequent silicic eruptions. This pattern coincides with pre-eruptive crustal uplift driven by mantle plume impingement. We propose that CO consumption associated with enhanced erosion and weathering of kilometer-thick Yangtze craton carbonates from regional uplift might have temporarily exceeded the CO contribution from magmatic degassing. This is supported by geochemical data indicating that the Emeishan basalt is particularly CO -poor. Our findings demonstrate that LIP environmental impact begins before the main eruptive phase, highlighting that tectonic-magmatic interactions can produce more complex CO patterns than previously recognized. This may explain why some LIPs caused extinctions, whereas others did not.

Introduction
Large igneous provinces (LIPs) represent the most extensive volcanic events in Earth’s history, typically outpouring over 100,000 cubic kilometers of lava within a short duration of 1 to 5 million years1. The conventional model links LIP volcanism to massive carbon dioxide (CO ) emissions that drive global warming and environmental perturbations2,3. However, recent studies reveal that the relationship between LIP emplacement and atmospheric CO is more complex. In the Central Atlantic Magmatic Province, episodic eruptions produce immediate CO spikes, followed by declines attributed to enhanced weathering of freshly exposed volcanic rocks4. However, in the case of the Deccan Traps and Columbia River Basalt Group, CO emissions are observed to increase prior to the onset of their main flood basalt volcanism due to the pre-eruption intrusions5. Moreover, “cryptic” carbon emissions can persist for several million years after the main eruptive phase through sustained magmatic degassing and metamorphism of organic-rich sediments6. These varying patterns might reflect differences in magma volatile content, intrusive activity, crustal architecture, and the relative timing of competing processes, such as volcanic degassing and enhanced weathering.

The environmental impact of LIPs was particularly significant during the Permian Period, which experienced two distinct biotic crises. The end-Guadalupian event (~260 million years ago (Ma)) was a critical precursor that possibly set the stage for the devastating end-Permian mass extinction7,8. This event notably affected shallow-marine taxa such as fusulines, corals and brachiopods across the Guadalupian-Lopingian (G-L) boundary9–13. Geochronological data have established temporal synchronicity between the end-Guadalupian mass extinction and the emplacement of Emeishan LIP14–18. While volcanogenic CO has been hypothesized as the key driver of this biotic crisis19, the actual relationship between the Emeishan volcanism and atmospheric CO concentrations (pCO ) remains poorly constrained.

Here we present a high-resolution pCO record spanning the Mid to Late Permian using compound-specific carbon-isotope measurements of chlorophyll-derived biomarkers. By resolving pCO changes through different phases of the Emeishan LIP development, we evaluate whether this volcanic episode conforms to the conventional LIP–CO rise–extinction model, and how carbon cycle dynamics during this interval may have influenced the end-Guadalupian biotic crisis.

Results
High-resolution stable carbon isotope records of bulk sediments (δ13C , δ13C ) and biomarker phytane (δ13C ) (Fig. 1) from exposed strata at Shangsi section (32°20′N, 105°28′E, Supplementary Fig. 1) document well-preserved marine sedimentation spanning from the Guadalupian (including the Roadian, Wordian, and Capitanian stages) to the Lopingian. One characteristic signature of G-L boundary sections is the negative carbon isotope excursion (CIE) recorded in carbonates across various regions (Supplementary Fig. 2), although localized depositional hiatuses may obscure this trend. Our δ13C profile shows a gradual negative shift from ~4‰ to a minimum of ~1‰ during the transition from the Guadalupian to the Lopingian (Fig. 1c), followed by a return to pre-CIE levels (average of 4‰) during the Wuchiapingian stage. However, the δ13C record exhibits a different pattern (Fig. 1d). At the G-L boundary, instead of a negative excursion, the values reach a transient maximum of ~−23‰. This is followed by two distinct negative shifts throughout the Wuchiapingian stage, each interrupted by a temporary positive excursion.

Phytane is a diagenetic product derived from the phytol side chain of chlorophyll produced by photoautotrophic organisms, primarily algae and cyanobacteria20. The carbon isotopic composition of phytane preserves information about isotopic fractionation that occurred during photosynthetic carbon fixation by the primary producers that synthesized the original chlorophyll molecules. Consequently, the compound-specific δ13C values of sedimentary phytane can serve as an effective proxy for reconstructing geological pCO levels when appropriately calibrated21–24. At Shangsi, our δ13C measurements show a positive shift from ~−31‰ to −27‰ from the Capitanian into the Early Wuchiapingian, followed by a gradual decrease to a minimum of ~−32‰ (Fig. 1e). When converting δ13C values to the original biomass δ13C, we applied the established 3.3‰ enrichment factor documented by ref. 22. This correction accounts for the systematic isotopic fractionation that occurs during chlorophyll biosynthesis and subsequent diagenetic transformation to phytane.

To reconstruct atmospheric CO levels, we quantified the isotopic fractionation (ε ) associated with carbon fixation during photosynthesis using our paired values of δ13C and δ13C following established methods25 (See Methods). This parameter increases systematically with rising pCO , though it can also be influenced by factors such as phytoplankton growth rate, cell geometry, and active carbon transport mechanisms26–28. In Fig. 1f, the calculated ε profile shows a negative shift from ~22‰ to a minimum of ~17‰ during the transition from the Capitanian to the onset of the Wuchiapingian, followed by a pronounced increase in ε . By the mid-Wuchiapingian stage, ε reaches a maximum value of ~24‰, and subsequently decreases to ~21‰.

New geochronological data provide a precise temporal framework of the eruption sequence within the Emeishan volcanic province14–17. The main eruptive phase of the Emeishan flood basalt volcanism occurred from 260.55 Ma to 259.1 Ma (refs. 15,17), with the earliest documented eruption at ca. 263.5 Ma (ref. 14). The waning stage of Emeishan LIP volcanism extended until ca. 257.4 Ma in the early Wuchiapingian16.

Our ε -based estimates of pCO (Fig. 2e) show that CO concentrations remain stable at approximately 700 parts per million by volume (ppm) throughout the Roadian and Wordian stages. Coincident with the onset of the Capitanian stage, pCO values begin to decrease at ~263.5 Ma, reaching an estimated minimum of 369 ppm, with a running mean average of ~550 ppm across the G-L boundary. This interval of decreasing pCO , from ~263.5 to 259.1 Ma, temporally corresponds with the early and main phases of the Emeishan flood basalt volcanism (Fig. 2a,e). Counterintuitively, the minimum pCO level coincides with the termination of the main phase of flood basalt eruptions, after which pCO values increase despite reduced eruption volumes. This surprising relationship is further highlighted by the correspondence between lowest pCO values and peak mercury (Hg) anomalies19 (Fig. 2b), demonstrating that the most intense volcanic activity during the main eruption phase paradoxically resulted in atmospheric CO drawdown rather than the expected increase. In the Early Wuchiapingian, pCO value reaches an estimated maximum of 1000 ppm, with a running mean average of ~900 ppm. This interval of elevated CO level coincides with explosive silicic eruptions rather than the earlier flood basalt volcanism16,29. Following the termination of the Emeishan volcanism, the pCO values decrease to approximately 600 ppm.

Our observed pCO trend from the Emeishan LIP reveals several unexpected findings in carbon cycle behavior that fundamentally challenge conventional understanding of LIP-climate interactions. First, while the standard model predicts atmospheric CO increases during peak flood basalt volcanism4, the Emeishan LIP exhibits precisely the opposite trend: a sustained CO decrease from ~700 to ~350 ppm during its most intense volcanic phase. This decoupling between volcanic activity and atmospheric CO is further highlighted by the correspondence between minimum pCO values and peak Hg anomalies, demonstrating that the most vigorous eruption phase paradoxically coincided with maximum carbon drawdown rather than addition. Second, our data reveal that the negative CIE at the G-L boundary occurred alongside CO drawdown rather than increase, contrasting with the typical coupling between negative CIEs and CO spikes observed in numerous other environmental perturbations throughout Earth history23,30,31. Third, the subsequent CO increase during the later silicic eruptive phase—despite its lower eruption volume—further demonstrates that eruption intensity does not directly correlate with atmospheric CO response in this case, reversing the expected pattern of volcanogenic carbon forcing. These multiple discrepancies from established patterns collectively indicate that a distinct and previously unrecognized mechanism must have dominated the carbon cycle during the Emeishan event, effectively counterbalancing and even overwhelming the expected volcanic CO contributions.

Methods
Materials
Samples were collected from the Shangsi section (32°20′N, 105°28′E, Supplementary Fig. 1), Shangsi Village, Guangyuan City, Sichuan Province, China. The Shangsi section, a prominent candidate for the Global Stratotype Section and Point (GSSP) of the Permian-Triassic boundary, is distinguished by its well-exposed marine successions and extensive research foundation. Based on paleogeographic reconstruction, Shangsi was located in the eastern region of the Paleo-Tethys during the Middle to Late Permian. All samples were collected from outcrops where the surface layers had been completely removed, exposing fresh layers to avoid weathered or otherwise contaminated rocks. Stratigraphically, the Shangsi section includes the Maokou and Wuchiaping formations, which correspond to the Guadalupian and Lopingian series, respectively. The G-L boundary is placed at the base boundary of the Wuchiaping Formation56. Fossil records indicate that the end-Guadalupian extinction was a multi-phase biotic crisis characterized by a gradual decline in biodiversity and episodic turnovers beginning in the mid-Capitanian. It affected multiple taxonomic groups such as fusulines, corals, ammonoids, brachiopods, conodonts and bivalves, leading to widespread ecosystem collapse9–13.

Chronology and Intrusive activity
The age control for the G-L interval is derived from the latest high-precision geochronology. The tie-points are as follows: 273.01±0.14 Ma for the base of the Roadian57, 266.9±0.4 Ma for the base of the Wordian58, 264.28±0.16 Ma for the base of the Capitanian58, 259.51±0.21 Ma for the base of the Wuchiapingian59 and 254.14±0.07 Ma for the base of the Changhsingian60.

The Emeishan intrusive system is expressed by a cluster of layered mafic-ultramafic complexes in the inner zone, typically surrounding the Panxi area (Panzhihua-Xichang, Supplementary Fig. 3a), including Panzhihua, Baima, Hongge, Taihe, Binggu and Xinjie18,52–54,61. Geophysical imaging along a seismic profile crossing the Emeishan LIP shows that crustal thickening and magmatic modification are most pronounced in the inner zone, where a thick mafic underplating layer is present, whereas the intermediate and outer zones lack comparable underplating features, implying that magmatic intrusion in these zones was relatively limited62. To date, only a single field study has documented composite, zoned diabase sills occurring in the outer zone, at Luodian in southern Guizhou63. Overall, the prevailing view on the distribution of intrusions in the Emeishan LIP is that they are mainly concentrated in the inner zone.

The timing of the Emeishan LIP intrusions broadly overlaps the main basaltic eruptions and extends from the earliest magmatic records to the waning stage (Supplementary Fig. 3d; refs. 14–18,52–54,61,63,64). Although the current consensus is that Emeishan LIP extrusive and intrusive activities were essentially coeval, the analytical quality and precision of available age determinations for the intrusions are insufficient to establish a precise intrusive chronology comparable to that of the Siberian Traps.

The country rocks intruded by the Emeishan bodies are predominantly carbonates. In the inner zone, intrusions were emplaced largely into Neoproterozoic (Sinian) dolostones that are nearly pure (containing very little clay or quartz), interbedded with siliceous limestones, marlstones and minor shales of the Dengying Formation65. In the outer zone, the Luodian sills intrude Early-Middle Permian Sidazhai Formation carbonates63. The development of contact aureoles within these carbonate successions suggests metamorphic decarbonation and localized carbonate partial melting of the country rocks during emplacement65,66.

Sample cleaning and preparation
All rock samples were cleaned and prepared following procedures described in ref. 67. Specifically, the rock samples were sonicated in three solvents (methanol, dichloromethane, and n-hexane) to remove any organic contamination. After cleaning, the samples were crushed into fine powder using a ring and puck mill and then stored in pre-combusted glass jars.

Bulk isotopic analysis
Bulk carbon isotopic values of carbonate (δ13C ) were analyzed using a Thermo Scientific Gas Bench coupled with a Delta V Plus isotope ratio mass spectrometer (IRMS), with Chinese national standards, GBW 04416 and GBW 04417 as references, at the Institute of Geology and Geophysics, Chinese Academy of Sciences. The carbon isotopic ratios are reported using standard δ notation (δ13C in ‰) relative to Vienna Pee Dee Belemnite (VPDB). Standard precision is better than 0.05‰ (1σ).

Rock powders were completely decarbonated by reacting with dilute hydrochloric acid and subsequently rinsed with Barnstead Nanopure water until the pH was neutral. Bulk carbon isotopic values of organic matter (δ13C ) were analyzed using a Thermo Scientific Flash elemental analyzer (EA) coupled with a Delta V Plus IRMS, with in-house laboratory standards along with reference standards IAEA-600 and USGS 40, at the Institute of Geology and Geophysics, Chinese Academy of Sciences. All samples were measured in triplicate. Standard precision is better than 0.1‰ (1σ).

Lipid extraction and biomarker analysis
Samples were extracted using a Thermo Scientific™ Dionex™ ASE™ 350 Accelerated Solvent Extractor in 9:1 dichloromethane/methanol mixture at 100°C to obtain total lipid extracts (TLEs). To maximize biomarker yield, more than 150 g of powdered rock was used per sample for extraction, and three ASE extraction cycles were performed until the extract color became lighter, ensuring sufficient biomarker abundance for subsequent isotopic analyses. Biomarker analysis was conducted at the Stable Isotope Geoscience Facility, Texas A&M University and the Institute of Urban Environment, Chinese Academy of Sciences. Saturated hydrocarbons were eluted through SiO -gel chromatography with 100% hexane. To remove n-alkanes, these fractions were transferred to cyclohexane and treated with activated 5Ǻ molecular sieves at 85°C for 24 h. Before isotopic analysis, we verified phytane identity and purity by GC-MS and required baseline-resolved phytane peaks with good peak shape. Any chromatogram showing peak overlap (coelution) was excluded. Representative GC-MS chromatograms illustrating phytane identification and baseline separation are shown in Supplementary Fig. 7. Compound-specific isotopic compositions of phytane (δ13C ) were analyzed using a Thermo Trace Ultra gas chromatograph coupled with a MAT 253 IRMS, equipped with a HB-5 column (50 m × 0.2 mm × 0.33 μm) and a programmable temperature vaporization injector with helium as the carrier gas. The oven program was set at 60°C upon sample injection, held for 1.4 min, then ramped up uniformly at 10°C min−1 to 180°C, 4°C min−1 to 260°C, and 20°C min−1 to 320°C, where it was held for 10 min. An external standard of n-alkanes with known δ13C values (A7, Arndt Schimmelmann, Indiana University) was used to monitor the instrument’s performance. The carbon isotopic ratios are reported relative to Vienna Pee Dee Belemnite. Hexadecane (C ) or eicosane (C ) with known δ13C values was used as the internal standard. All samples were measured in duplicates and corrected for drift and linearity. Standard precision is better than 0.3‰.

Terrestrial organic matter input. To assess the contribution of terrigenous organic matter in the Shangsi samples, we used n-alkane and isoprenoid indices (Supplementary Table 2). The abundances of these compounds were determined using an Agilent 8890 gas chromatograph equipped with a flame ionization detector (GC-FID) and a DB-5 capillary column (30 m × 0.25 mm × 0.25 μm). The temperature program started at 70°C, then increased at a rate of 10°C/min to 210°C, followed by 3°C/min to 300°C, where it was held for 30 min. Based on peak areas, we calculated (i) the terrigenous-aquatic ratio (TAR; TAR = [C +C +C ]/[C +C +C ]), which contrasts long-chain (C –C ) higher-plant wax alkanes with short-chain (C –C ) algal/bacterial alkanes68, (ii) the carbon preference index (CPI; CPI = [C +C +C +C ]/[C +C +C +C ]) of long-chain n-alkanes as an indicator of odd-over-even predominance typical of fresh higher-plant waxes69,70, and (iii) pristane/phytane (Pr/Ph), which generally reflects redox and depositional setting71. At Shangsi, 33 out of 39 samples have TAR values below 1, with an average of 0.58 for these samples. The six samples with TAR values greater than 1 are not concentrated at any particular stratigraphic level but occur randomly throughout the section. Long-chain CPI values cluster around 1 (0.67–0.98; average = 0.89±0.07, n=39), and Pr/Ph ranges from 0.23–1.03 (average = 0.59±0.14; n=38), with no values exceeding 1.5. Together, these indices indicate that the organic matter in the Shangsi samples is primarily derived from aquatic/marine sources, with a relatively minor higher-plant contribution.

pCO calculations
Phytane-based pCO reconstructions rely on the measurements of carbon isotopic composition from both phytane (δ13C ) and associated carbonates (δ13C ). The carbon isotope fractionations (ε ) during phytoplankton photosynthesis are utilized to reconstruct seawater and atmospheric CO concentrations throughout the Phanerozoic era. ε values are derived from the δ13C values of dissolved CO (δ ) and photosynthetic biomass (δ )(ref. 25). δ is calculated from the isotopic difference between phytol and photoautotrophic biomass. Prior culture studies estimate this difference to be 3.3±1.3‰ (ref. 22). Meanwhile, δ is calculated from δ13C , using the temperature-dependent relationship between the δ13C values of dissolved CO and solid carbonate72,73. The coeval seawater temperatures were derived from a compilation of high-resolution conodont apatite δ18O records57,74,75. Given the varying sampling resolutions, a LOESS smoothing function was applied to derive continuous temperature data across the timeline. The 1σ (68%) and 2σ (95%) uncertainties in ε calculations were determined through Monte Carlo simulations that accounted for various sources of uncertainty, including analytical uncertainty in δ13C , a standard deviation (SD) of 1.3‰ in the isotopic offset between phytane and biomass, and a 4°C SD in seawater temperature estimation.

In pCO reconstructions, it is assumed that dissolved CO is primarily transported to the site of carbon fixation in photosynthetic cells by diffusion76. Carbon concentrating mechanisms (CCMs) were unlikely to be fully active during the relatively high CO levels of the Permian and Triassic periods22, as this active carbon uptake by CCMs requires substantial energy, a characteristic observed in many modern algae species77. The concentrations of aqueous CO [CO (aq)] were calculated using the equation: ε = ε − b/[CO (aq)], where ε represents the net kinetic isotope fractionation of cellular enzymatic processes, b represents the physiological factors78–80. In this study, ε was assigned a value of 28‰. The b-value was calculated using the equation: b=82.05×[PO 3−]+86.28, where the [PO 3−] represents a phosphate concentration of 0.2 μM as a representative level. Further discussion on ε and b-value are discussed below. Finally, atmospheric pCO concentrations were calculated from CO (aq) using Henry’s law under the assumption of air-sea equilibrium81. The 1σ (68%) and 2σ (95%) uncertainties in pCO calculations were determined through Monte Carlo simulations that accounted for various sources of uncertainty, including analytical uncertainty in δ13C , a SD of 1.3‰ in the isotopic offset between phytane and biomass, a 4°C SD in seawater temperature estimation, and uncertainty in [PO 3−], which ranged from 0.17 to 0.27 μM.

Uncertainties in pCO estimates
The calculation of pCO involves various assumptions, including the dependence on [PO 3−] and the value of enzymatic isotope fractionation (ε), along with potential uncertainties in seawater temperature that can impact pCO estimates. These factors are discussed below.

Estimates of b. In the process of photosynthetic carbon fixation, the parameter b synthesizes the effects of growth rate, cellular morphology, and membrane permeability on carbon isotopic discrimination. In this paper, we adopt the estimates for b utilized in previous CO reconstructions for the Permian-Triassic boundary23. The refined equation for b as a function of [PO 3−] is derived from sediment data and is expressed as b=82.05 × [PO 3−]+86.28 (ref. 82). This formula yields b values that enhance the accuracy of CO reconstructions, providing better alignment with sedimentary records. In this equation, we used a representative [PO 3−] concentration of 0.2 μM at a depth of 20 m, representing the average value within the modern observed range of 0.17 to 0.27 μM from the ocean surface to a depth of 50 meters. Due to the absence of data on the temporal variability of b, we maintained it as a constant value throughout the study interval. Regardless of the uncertainty of b, its impact is confined to the absolute values of pCO , as sensitivity analysis demonstrates that pCO changes by less than 6% when the maximum or minimum b values are applied (see Supplementary Fig. 4). Consequently, these variations do not alter the overall trend of pCO calculations over time.

Estimates of ε. Enzymatic isotope fractionation (ε) is primarily driven by the activity of the enzyme Rubisco and/or contributions from β-carboxylases. Experimental studies on the growth of eukaryoticalgae and cyanobacteria have determined that ε values typically are within the range of 25‰ to 28‰ (refs. 27,79,83,84). In this study, we adopt the ε value of 28‰, as β-carboxylation is typically inhibited under higher CO levels, leading to larger ε value85. Supplementary Fig. 5 shows the theoretical relationship among ε , phosphate concentration, and seawater temperature, based on the ε value of 28‰.

A recent modern-ocean calibration of the phytol ε -[CO (aq)] relationship, based on a Deming regression, reports ε ≈23.9‰ and b≈82.4 (ref. 24). This study refines and strengthens the phytol/phytane-based approach to pCO reconstruction, thereby improving its reliability and expanding its applicability for future studies. In our study, the b value is parameterized from the phosphate relationship, which yields b≈102. This value is substantially lower than the traditional constant of b=170 and broadly consistent with the low-b range in modern calibrations24. As for ε, while the modern calibration provides an important reference framework, several considerations suggest that its ε value may not be directly applicable to our study. First, the calibration domain of modern datasets is limited at the high-[CO (aq)] end. The open ocean rarely attains very high [CO (aq)], and near-shore waters, where [CO (aq)] can be higher, are typically excluded to minimize terrestrial inputs. This restriction reduces the opportunity to observe maximum fractionation, thereby leading to a lower fitted ε (ref. 24). Second, carbon-concentrating mechanisms (CCMs) and active HCO -uptake (β-carboxylation) are pervasive in modern phytoplankton and tend to depress the apparent fractionation. During deep-time intervals, these pathways were likely down-regulated under elevated pCO , implying a higher effective ε. In addition, because of the asymptotic nature of the ε -pCO relationship, adopting a low ε places more of our data near the ε ≈ε limit, where small uncertainties in ε can propagate into disproportionately large pCO (ref. 24). For these mechanistic and methodological reasons, we adopt ε =28‰ in our pCO reconstructions.

Correlation of records of δ13C
The G-L CIE is characterized by a pronounced decrease in δ13C values near the boundary57. Our δ13C profile exhibits trends and patterns that are consistent with those of other G-L records observed in various paleo-continental regions (Supplementary Fig. 2). This consistency suggests that the δ13C record from Shangsi reflects the typical marine δ13C signal, which was subsequently used to calculate the isotopic composition of dissolved CO .

LOESS
The running average was generated using the locally estimated scatterplot smoothing (LOESS) function in SigmaPlot V14 software. For this calculation, a first-order polynomial was applied as the smoothing degree, and the sampling frequency was set to 0.4 to define the proportion of data points included in the smoothing process.

Estimating the impact of Maokou Formation weathering on the carbon cycle
Estimating the volume and mass of eroded Maokou limestone. The uplifted area of the Maokou Formation is approximately 2,010,619 km2, assuming a circular area with a radius of 800 km (ref. 32). The estimated average erosion thickness is 150 m (ref. 32), resulting in an eroded volume of 301,593 km3 (calculated as 2,010,619 km2 × 0.15 km). With an average limestone density of 2.5 × 1012 kg/km3, the total mass of the eroded limestone is calculated to be 7.54 × 1017 kg or 754,000 Gt.

Total organic carbon content in Maokou Formation samples. To evaluate the total organic carbon (TOC) content of the Maokou Formation, we compiled previously published TOC data from different sections. Reported average TOC values are 1.64% at Chaotian86, 1.96% at Shangsi87, 7.58% at Xibeixiang88 and 0.19% at Tieqiao89. As the Xibeixiang data are limited to the uppermost Maokou Formation, they are excluded from subsequent analyses. The remaining three sections exhibit relatively consistent values, with a weighted average of 1.26% for the Maokou Formation. Note that the Shangsi value is taken from an independent study87 rather than our own measurements; however, our TOC data yield a similar value (2.13%), supporting the reliability of the published result.

Limestones generally contain less organic carbon than fine-grained siliciclastic rocks (e.g., shales) deposited in the same settings. The TOC content in most marine limestones is relatively low, typically below 0.5%. Modern carbonate sediments can accumulate organic matter in amounts comparable to shales, but much of the organic carbon is lost during burial through oxidation or decay, resulting in only a small residual TOC. However, certain limestone units do display elevated TOC under some specific conditions, such as anoxic bottom waters, high primary productivity and rapid burial, which facilitate the formation of organic-rich limestones.

Widespread anoxic conditions during the deposition of the Maokou Formation have been documented across various regions, including mid- to high-latitude areas of Panthalassa and the Tethys Ocean86,90–94. These oxygen-deficient conditions were not restricted to deep marine basins but also extended into intermediate to shallow water depths40,86,90–92,95. For instance, sulfur (δ34S) isotope records coupled with pyrite iron speciation analyses from the shelf carbonates at Chaotian reveal that the expansion of the oxygen minimum zone (OMZ) during the Capitanian drove the development of euxinic conditions86. Similarly, multiple sulfur isotopes (δ34S and Δ33S) in pyrites from the Penglaitan and Tieqiao sections further support the widespread shoaling of sulfidic waters90. Furthermore, negative shifts in δ238U values of marine limestones at Xikou also suggest intensified marine anoxia, with an estimated ~10% anoxic seafloor coverage lasting >3 Myr (ref. 96). In addition, remarkably high nitrogen isotope values (δ15N >6‰, with peaks ≥10‰) in Maokou Formation limestones have been reported across several regions, exceeding most Phanerozoic δ15N records86,91,97–99. In the modern deep ocean, nitrate δ15N values typically range from 4‰ to 6‰. However, strongly denitrifying OMZs (e.g., the eastern tropical Pacific Ocean and the Arabian Sea) often exhibit exceptionally high δ15N values of 15–20‰ (refs. 100,101). These regions with such high δ15N values are characterized by high-productivity and low-oxygen levels, where intense water-column denitrification occurs. Therefore, these findings suggest that the Maokou Formation was likely deposited under oxygen-depleted conditions, which may have facilitated the accumulation of relatively high TOC in carbonate sediments.

Estimating the carbon content of eroded material. Inorganic carbon (carbonate): assuming a carbon content of 12% in CaCO , the mass of inorganic carbon is estimated at 90,480 Gt C (calculated as 754,000 Gt × 0.12). Organic carbon: the average total organic carbon content is 1.26 wt%, resulting in an estimated mass of organic carbon of 9,500.4 Gt C (calculated as 754,000 Gt × 0.0126). The net effect on the carbon cycle yields a net consumption of 80,979.6 Gt C. Over a period of 3 million years, this results in a rate of change of approximately 0.027 Gt C per year.

Impact on the marine carbonate δ13C. Assuming the δ13C of weathered and original marine carbonate is approximately 0‰, and that of organic carbon is −25‰, the mass balance calculation for δ13C of the total input gives δ13C as (90,480 × 0‰ + 9500.4 × −25‰) / 99,980.4 = −2.375‰. Without buffering effects, the initial ocean carbon stands at 38,000 Gt C with a δ13C of +1‰. After including the input from weathering, which totals 99,980.4 Gt C at −2.375‰, the new δ13C calculates to (38,000 × 1‰ + 99,980.4 × −2.375‰) / 137,980.4 = −1.39‰. This suggests a potential negative shift in marine carbonate δ13C of approximately 2.39‰, from +1‰ to −1.39‰, over a period of 3 million years.

Considering buffering, the carbonate compensation assumes that for every mole of CO added from weathering, about 0.8 moles of carbonate sediment dissolve, totaling 99,980.4 Gt C × 0.8 = 79,984.32 Gt C. Assuming δ13C of carbonate sediments similar to the initial ocean DIC at +1‰, a new mass balance calculation for the total carbon, which now amounts to 38,000+99,980.4+79,984.32 = 217,964.72 Gt C, yields a new δ13C of (38,000 × 1‰ + 99,980.4 × −2.375‰ + 79,984.32 × 1‰) / 217,964.72 = −0.28‰. Thus, the buffered change in δ13C equates to −0.28‰ from the initial +1‰, resulting in a change of −1.28‰.

Impact on the atmospheric CO . The initial Permian atmospheric CO concentration was set to be 700 ppm. To convert this value into gigatons of carbon (Gt C), we assume the present-day atmospheric mass of 5.1 × 1018 kg and a molecular weight ratio of CO to air of 1.52. Using these values, 700 ppm is equivalent to 700 × 5.1 × 1018 × 1.52 / (106 × 106) = 5426.4 Gt CO . Converting this to carbon, we yield 5426.4 Gt CO × (12/44) = 1479.9 Gt C. Without buffering effects, the net consumption of 80,979.6 Gt C would far exceed the initial atmospheric CO content of 1479.9 Gt C, indicating the potential for complete depletion of atmospheric CO . However, considering buffering mechanisms such as carbonate compensation, we conservatively estimate that only 10% of the net change would affect the atmosphere. This results in 80,979.6 Gt C × 0.1 = 8097.96 Gt C net removal from the atmosphere, which still exceeds the initial atmospheric CO content. A more realistic scenario might involve atmospheric CO decreasing by approximately half, leading to a 1479.9 Gt C × 0.5 = 739.95 Gt C decrease. This would reduce the new atmospheric CO level to 1479.9 − 739.95 = 739.95 Gt C. Converting this back to ppm, we find that 739.95 Gt C is equivalent to 2713.15 Gt CO . Using the same conversion factors, we estimate the new atmospheric CO concentration to be approximately 350 ppm.

Sensitivity analysis of net CO consumption. To assess robustness, we quantified how the carbonate-weathering sink scales with the geometry of the uplifted dome and the organic-carbon content by exploring dome radius of 600–900 km, average erosion thickness of 50–300 m, and bulk TOC values of 0.5–2 wt%. For all combinations, the net carbon sink ranges from 1.41 × 104 Gt C at (600 km, 50 m, 2%) to 2.19 × 105 Gt C at (900 km, 300 m, 0.5%), and is therefore ≫ 103 Gt C under all tested conditions (Supplementary Fig. 6). Assuming a 2–4 Myr weathering interval, mean removal rates span 0.003–0.11 Gt C yr-1. Note that our estimate of the carbonate-weathering volume and rate represents a simplified, first-order approximation intended to illustrate the likely range of values rather than capture the full complexity of the system.

Peer review rounds
[{"round": 1, "reviewer_comments": {"reviewer_1": "Jiaheng Shen and colleagues present an impressive and thought-provoking study linking uplift associated with the\nemplacement of the Emeishan LIP to major disturbances to the surficial carbon cycle in the Permian. The measurements of\nvarious and contrasting excursions in carbon pools is a powerful tool for understanding environmental perturbations. Overall,\nI believe the study is worthy of publication in Nature Communications after minor to moderate revisions, which are detailed\nbelow.\nMain points:\nThe main issue I have with the current manuscript is the lack of subdivision in degassing sources during LIP emplacements.\nFor example, the statement in lines 34-35 “Unlike other LIPs where magmatic CO2 release dominates the atmospheric\nsignal” is not borne out by modern studies of LIPs. Magmatic degassing is a volumetrically important component of LIP\ncarbon, but is not believed to be rapid enough to cause mass extinction events alone (e.g., Jones et al., 2016). In contrast,\nthe rapid and widespread intrusions in the shallow crust can be a more volumetrically important, and faster, source of carbon\nto the atmosphere. This thermogenic source has been suggested as the trigger for the main climatic disturbance where sills\nwere emplaced in organic-rich sediments for the Siberian Traps (Burgess et al., 2017), the Central Atlantic Magmatic\nProvince (Heimdal et al., 2020), the Deccan Traps (Tian & Buck 2022 [ref. 5]), and the\nNorth Atlantic Igneous Province\n(Svensen et al., 2004; Jones et al., 2019). In several of these instances, the main effusive basalt volcanism predates the\nenvironmental disturbances, effectively acting as a cap that then favours intrusions. The impact of thermogenic degassing is\nheavily dependent on the extent of intrusive activity and the organic content and chemistry of country rocks (Svensen et al.,\n2023). I am not familiar with the extent of intrusive activity associated with the Emeishan LIP, but it is likely to be: a) a\nsignificant proportion of the total emplacement, and b) not necessarily synchronous with the main effusive lava\nemplacement. Including a more nuanced description of LIP degassing (expanding on lines 137-146) would help inform on\nthe surface carbon cycle signals presented in this manuscript.\nAnother potential issue is that the authors are considering the effects of the Emeishan LIP to be the primary cause of what\nmay be more global changes. This may be the case, but there are other potential candidates in the Permian for large-scale\nchanges in the carbon cycle. The Emeishan LIP occurs at the end of the Late Paleozoic Ice Age (LPIA), which recent studies\nhave suggested may have been instigated and ended by the availability of land within the tropical high-weatherability zones\nas Pangea moved slowly northwards (e.g., Goddéris et al., 2017; Marcilly et al., 2025). I think it is important to frame the\ncarbon cycle instabilities across the Capitanian-Wuchiapingian in a broader context of widespread carbon fluctuations\nacross the LPIA in the Carboniferous and Permian. Moreover, the placement of Emeishan within the high weatherability\nzones is important for the potential rate and extent of weatherability for a domed sequence. Including this perspective could\nenhance the author’s findings of rapid unroofing above a mantle plume.\nThe calculations on the rate and volume of carbonate weathering from uplifted strata seem a little simplistic. The combined\nsurface reservoirs of carbon in the Permian were much bigger than today, and the negative feedbacks that counteracted\npCO2 increase were different as there wasn’t a full-scale pelagic carbon pump set up yet (e.g., Ridgwell and Zeebe, 2005).\nA little more discussion is needed in how the unique forcings associated with Emeishan (location, uplift amount, size)\ninterplayed with the late Paleozoic carbon cycle in particular. This directly feeds into the catchy title of “Every LIP is\ndifferent”, and some more thorough investigation of why each LIP is different would be a welcome addition to the paper.\nMinor comments:\nLine 69: Change “biomarker” to “biomarkers”\nLine 182: Extensive uplift prior to the emplacement of a LIP is not a phenomenon unique to Emeishan. In particular, the\nNorth Atlantic Igneous Province (NAIP) has substantial uplift preceding the main phase of emplacement (Maclennan &\nJones, 2006; Conway-Jones & White 2022). The radius and uplift area for Emeishan are considerably less than calculated\nfor the NAIP, which may be indirect evidence of a smaller LIP volume. Therefore, the location of the uplift (i.e., whether it is in\na high weatherability zone) is going to be critical for supporting the conclusion that rapid carbonate weathering led to intense\ncarbon release.\nLine 234: A value of 135 ppm is way below CO2 concentrations in all modern systems. This suggests a preservation issue\nrather than a true representation of magmatic CO2 degassing. As mentioned in the main point, it is thermogenic degassing\nthat is likely to be the most important source.\nFigure 3: The schematic drawing of the LIP has a stratovolcano in the centre of the LIP, which is not accurate. I suggest\nredrawing to show Hawaiian/Icelandic type eruptions from fissures or a shield volcano.\nLine 296: This statement needs revisiting through the lens of when and how much of the Emeishan LIP was intrusive, and\nwhether there was significant thermogenic degassing associated with it.", "reviewer_2": "Key Message: This paper used a phytane-based phytoplankton proxy to reconstruct pCO2 values across the Emishan LIP\nevent during the Permian. The authors found that pCO2 varied during different stages of the LIP, rather than follow a general\nincrease, as is the expected norm during a LIP. Specifically, pCO2 decreased during the initial and main stage of the LIP\nand only increased in the final stage. Significant increases in chemical weathering, associated with uplift, is suggested as\nthe cause of the initial decrease.\nSignificance: This paper provides detailed d13C_carb, d13C_org, and d13C_phytane measurements, the latter of which\nhave a very limited published record in the Permian. Filling in these gaps has enabled the authors to reconstruct pCO2 over\nthis time period using the phytoplankton pCO2 proxy based on phytane. While this would be significant on its own, the\nauthors also discovered an unexpected decrease in the pCO2 values across the Emishan LIP, providing nuance to our\nunderstanding of how LIPS influenced the carbon cycle.\nData/Methods: The authors used d13C_carb, d13C_org, and d13C_phytane to calculate the amount of carbon fractionation\noccurring in phytoplankton (epsilon_p). This value is then used to calculate pCO2 using previously established methods as\nmentioned in the paper. The data and methods are sound, however there are some potential gaps in determining the quality\nof the data, specifically:\n1) Given the very low TOC values of many of the samples, how well was phytane isolated from the background signal?\n2) Was there any evidence of terrestrial input in the samples that may have altered the organic content?\nIn addition, Graham et al. (2025) recently calibrated epsilon_p of phytol to pCO2 in modern sediments and applied their\ncalibration to phanerozoic data sets, including phytane data sets. The authors should consider utilizing this calibration or\naddress why they have chosen not to use it.\nClarity/context: The paper is well written and delivers its key messages clearly. Figs 1 and 2 are easy to interpret and do a\ngood job of condensing and comparing multiple lines of evidence. The small volcano in the bottom left corner of fig 2 is a bit\nconfusing and doesn’t match with the rest of the fig layout.\nFIG 3 - The order of fig 3 is also a little confusing as the bottom section is labeled ‘a’ and the top ‘d’. Most readers will read\nthe fig from top to bottom, and the labels should be switched to reflect this order, or both the figure and labels should be\nswitched (i.e. top = ‘a’ pre-mantle; bottom = ‘d’ stage III). A timeline attached to the side might also clarify the order in which\nthe figures appear, so that the audience immediately knows which are the older stages and which are the younger.\nThe text in this figure is also very small and it is unclear why some text is distinguished by colored boxes behind it and some\nis not.\nIt is also not clear why Stage III specifies composite weathering input ‘dominant crust-derived carbon’ while Stage I and II\nseparate organic and carbonate weathering inputs. The relevance of the thickness of the carbon arrows should be noted in\nthe figure, not just the description.\nThe dashed line mentioned in the text description is orange, not red.\nExpertise: I have kept most of my comments related to the organic geochemistry and carbon-isotopes section of this paper\nas that falls within my expertise. I do not have strong expertise in Permian LIPs and mantle dynamics and as such have a\nlimited assessment of those sections."}, "author_response": "We would like to express our sincere gratitude to the editor and the reviewers for their\nvaluable comments and constructive suggestions. All points raised by Reviewer #1 and\n#2 are addressed in detail below (in blue) and revised accordingly in the manuscript.\nReviewer #1 (Remarks to the Author):\nJiaheng Shen and colleagues present an impressive and thought-provoking study\nlinking uplift associated with the emplacement of the Emeishan LIP to major\ndisturbances to the surficial carbon cycle in the Permian. The measurements of various\nand contrasting excursions in carbon pools is a powerful tool for understanding\nenvironmental perturbations. Overall, I believe the study is worthy of publication in\nNature Communications after minor to moderate revisions, which are detailed below.\nWe thank the reviewer for the positive comment. Below, we addressed the reviewer’s\nconcerns point by point.\nMain points:\nThe main issue I have with the current manuscript is the lack of subdivision in degassing\nsources during LIP emplacements. For example, the statement in lines 34-35 “Unlike\nother LIPs where magmatic CO2 release dominates the atmospheric signal” is not borne\nout by modern studies of LIPs. Magmatic degassing is a volumetrically important\ncomponent of LIP carbon, but is not believed to be rapid enough to cause mass\nextinction events alone (e.g., Jones et al., 2016). In contrast, the rapid and widespread\nintrusions in the shallow crust can be a more volumetrically important, and faster,\nsource of carbon to the atmosphere. This thermogenic source has been suggested as the\ntrigger for the main climatic disturbance where sills were emplaced in organic-rich\nsediments for the Siberian Traps (Burgess et al., 2017), the Central Atlantic Magmatic\nProvince (Heimdal et al., 2020), the Deccan Traps (Tian & Buck 2022 [ref. 5]), and the\nNorth Atlantic Igneous Province (Svensen et al., 2004; Jones et al., 2019). In several of\nthese instances, the main effusive basalt volcanism predates the environmental\ndisturbances, effectively acting as a cap that then favours intrusions. The impact of\nthermogenic degassing is heavily dependent on the extent of intrusive activity and the\norganic content and chemistry of country rocks (Svensen et al., 2023). I am not familiar\nwith the extent of intrusive activity associated with the Emeishan LIP, but it is likely to\nbe: a) a significant proportion of the total emplacement, and b) not necessarily\nsynchronous with the main effusive lava emplacement. Including a more nuanced\ndescription of LIP degassing (expanding on lines 137-146) would help inform on the\nsurface carbon cycle signals presented in this manuscript.\nWe thank the reviewer for this constructive and important suggestion. We agree that\nLIP emissions include both magmatic and thermogenic components, and that rapid\nthermogenic release linked to sill emplacement in sedimentary basins can dominate\natmospheric signals and, in some cases, lead to severe environmental crises, such as the\nwell-known Siberian Traps. In the revision, we followed the reviewer’s suggestion by\n(i) providing a synthesis of intrusive/sill emplacement and its potential degassing\nsource for the Emeishan LIP and, building on this, (ii) including a brief comparison\nwith the Siberian Traps, as both LIPs occurred during the Permian and, in our previous\nwork (Shen et al., 2022), we reconstructed CO records across the P-Tr boundary. We\n2\nbelieve that this comparison helps to better highlight the main theme of the paper—that\nevery LIP is different. Note that the phrase “Every LIP is different:” has now been\nremoved from the title in order to comply with the journal’s title format (no punctuation\nor puns).\n(i) Intrusive activity during the Emeishan LIP emplacement\nMain intrusive types and distribution: For the Emeishan LIP, the intrusive system is\nexpressed by a cluster of layered mafic-ultramafic complexes/intrusions in the inner\nzone, typically surrounding the Panxi area (Panzhihua-Xichang, Fig. A-a), including\nPanzhihua, Baima, Hongge, Taihe, Binggu and Xinjie (Zhou et al., 2005; Zhong and\nZhu, 2006; Zhong et al., 2011; Shellnutt et al., 2011; Zhou et al., 2002). These intrusions\nare generally 10–20 km long (except for the Binggu and Xinjie intrusions, which are\n~7 km and ~2 km long, respectively) and up to ~2 km thick, forming sill-like to S-\nshaped bodies that host giant Fe-Ti-V oxide ore deposits (Zhou et al., 2005; Pêcher et\nal., 2013). Geophysical imaging along a seismic profile crossing the Emeishan LIP\nshows that crustal thickening and magmatic modification are most pronounced in the\ninner zone, where a thick mafic underplating layer is present, whereas the intermediate\nand outer zones lack comparable underplating features, implying that magmatic\nintrusion in these zones was relatively limited (Chen et al., 2015). To date, only a single\nfield study has documented composite, zoned diabase sills occurring in the outer zone,\nat Luodian in southern Guizhou (Huang et al., 2019). Overall, the prevailing view on\nthe distribution of intrusions in the Emeishan LIP is that they are mainly concentrated\nin the inner zone and that the overall extent of sill intrusion is less pervasive than in the\nSiberian Traps.\nThe timing of the Emeishan LIP intrusions broadly overlaps the main basaltic eruptions\nand extends from the earliest magmatic records to the waning stage (Fig. A-b). In the\ninner Panxi region, zircon U-Pb ages from the Baima, Hongge, Binggu, and Xinjie\nintrusions cluster around the main eruptive pulse (~261–258 Ma), with the majority of\ndates concentrated toward its termination (~259 Ma; Fig. A-b, yellow starred symbols).\nWithin the same region, the Panzhihua and Taihe complexes show a broader temporal\nrange, with the oldest ages approaching ~264 Ma and the youngest extending into the\npost-LIP stage (~255 Ma). In contrast, in the outer zone, sills are so far constrained by\nonly a single dated occurrence (260.0 ± 3.4 Ma), and the wide analytical uncertainty\ntogether with the limited number of dated samples hampers a clear assessment of their\nfull temporal range. Overall, the analytical quality and precision of currently available\nage determinations for the Emeishan LIP intrusions are insufficient to establish a\nprecise intrusive chronology comparable to that of the Siberian Traps.\nFig. A | Map and time framework of the Emeishan LIP. (a) Reported distribution of\nEmeishan flood basalts and mafic-ultramafic intrusions. (b) Published chronological\nrecord of Emeishan LIP magmatism based on Zircon U-Pb dates. Data from refs. (Zhou\net al., 2005; Zhong and Zhu, 2006; Zhong et al., 2011; Shellnutt et al., 2011; Zhou\net al., 2002; Huang et al., 2019; Tang et al., 2021).\nThe country rocks intruded by the Emeishan intrusive bodies are predominantly\ncarbonates. In the inner Panxi region, intrusions were emplaced largely into\nNeoproterozoic (Sinian) dolostones that are nearly pure (very little clay or quartz),\ninterbedded with siliceous limestones, marlstones and minor shales of the Dengying\nFormation (Ganino et al., 2013). A broad contact aureole developed in these carbonate\nsuccessions, dominated by brucite marbles with local calc‑silicate/skarn horizons and\nbanded reaction zones around mafic dykes (Ganino et al., 2008; 2013). These features\nindicate metamorphic decarbonation and local carbonate partial melting of the country\nrocks during emplacement (Ganino et al., 2013). In the outer zone, the Luodian sills\nintrude Early-Middle Permian Sidazhai Formation carbonates (Huang et al., 2019).\nGanino and Arndt (2009) propose that the primary degassing during Emeishan LIP\nemplacement was derived from thermogenic/metamorphic CO from carbonate country\n2\nrocks, rather than direct magmatic degassing.\n(ii) comparison with the Siberian Traps\nIn the Siberian Traps, widespread sills intruded evaporite- and organic-rich sediments\n(Cambrian evaporites and the Carboniferous-Permian coals) in the Tunguska Basin,\nwhich rendered the province particularly lethal and led to extreme atmospheric and\nenvironmental disruption (Svensen et al., 2023). High‑precision geochronology and\nbasin stratigraphy define three magmatic stages with distinct emplacement styles in\nSiberian LIP (Burgess and Bowring, 2015). A transition from dominantly extrusive\nlavas (Stage 1) to widespread sill intrusions marks the onset of Stage 2 (Burgess and\nBowring, 2015). During Stage 2, spatially extensive sill complexes emplaced across the\nvolatile-rich Tunguska basin, generating enormous quantities of thermogenic gases that\nwere rapidly vented to the atmosphere through hydrothermal pipes over a vast area\n(Svensen et al., 2009; 2018; 2023; Burgess et al., 2017). These studies identify the\nStage-2 shift as the critical driver of carbon release and climatic perturbation. Our\nprevious reconstructed CO record (Shen et al., 2022) independently support this\n2\nmechanism: no significant CO increase is observed during Stage 1, whereas a\n2\npronounced rise begins at ~251.9 Ma, synchronous with the onset of Stage 2 sill\nemplacement. These observations highlight the important role of the sill-triggered\nthermogenic degassing in the Siberian LIP’s catastrophic environmental consequences.\nIn contrast, comparable province-wide sill intrusion into organic-rich sedimentary\nsequences analogous to Tunguska Basin has not been demonstrated in the Emeishan\nLIP. The coal-bearing Longtan Formation, although rich in organic matter, generally\npost-dates the main phase of the Emeishan flood basalts, overlies them, and is not\nclearly intruded by sills of Emeishan-LIP age (Li et al., 2017; Qiu et al., 2019; Shen et\nal., 2021). Available geological observations suggest that metamorphic carbon derived\nfrom decarbonation of carbonate country rocks was likely the main degassing source\n(Ganino and Arndt, 2009). However, because the precision of existing age data for\nEmeishan sills and intrusions is limited, the relative contribution and timing of\nthermogenic CO release from these intrusions cannot be precisely constrained. Even\n2\nallowing for this carbon release, the restriction of intrusions to the inner zone implies\nthat the atmospheric CO signal was still dominated by weathering, since uplift and\n2\nsubsequent erosion of carbonates represent a much larger CO sink.\nnuanced description of the Emeishan LIP intrusion and degassing. Accordingly, we\nhave incorporated the above content into the revised manuscript and made the following\nchanges:\n• Removed the inaccurate phrase “Unlike other LIPs where…” from the Abstract.\n• Added a discussion in the main text (Lines 243–260).\n• Included additional details on the intrusions in the Methods (Lines 351–376).\n• Added the timing and distribution of the Emeishan intrusions in Supplementary Fig.3.\nOur revised text now reads:\nLines 243–260:\n“Beyond the magmatic volatile budget, thermogenic CO generated where sills intrude\n2\norganic-rich sedimentary basins can be both volumetrically significant and rapid46-50.\nIn the Siberian Traps, high‑precision geochronology documents a transition from early\nextrusive lavas/pyroclastics to widespread sill emplacement into the volatile‑rich\nTunguska Basin51. The emplacement of these spatially extensive sills led to enormous\nthermogenic degassing46-49. An independent pCO reconstruction23 shows a coeval CO\n2 2\nrise with the onset of this sill-dominated intrusive phase in the Siberian Traps,\nsupporting this interpretation. In contrast, the Emeishan LIP is characterized by layered\nmafic intrusions concentrated in the inner-zone Panxi region52-54 (see Methods for\ndiscussion of intrusive activity). Comparable province-wide sill intrusion into organic-\nrich sedimentary sequences analogous to those in the Tunguska Basin has not been\ndemonstrated in the Emeishan LIP. The predominance of carbonate host rocks and\nassociated contact aureoles around Emeishan intrusions suggest that metamorphic\ncarbon liberated by decarbonation of carbonate country rocks was likely the main\ndegassing source55. However, the limited precision of existing age data for Emeishan\nintrusions hinders precise constraints on the relative contribution and timing of\nthermogenic CO release. Nevertheless, because intrusions are confined to the inner\n2\nzone, their impact on atmospheric CO was likely limited.”\n“Intrusive activity during the Emeishan LIP emplacement\nThe Emeishan intrusive system is expressed by a cluster of layered mafic-ultramafic\ncomplexes in the inner zone, typically surrounding the Panxi area (Panzhihua-Xichang,\nSupplementary Fig. 3a), including Panzhihua, Baima, Hongge, Taihe, Binggu and\nXinjie18,52-54,61. Geophysical imaging along a seismic profile crossing the Emeishan LIP\nshows that crustal thickening and magmatic modification are most pronounced in the\ninner zone, where a thick mafic underplating layer is present, whereas the intermediate\nand outer zones lack comparable underplating features, implying that magmatic\nintrusion in these zones was relatively limited62. To date, only a single field study has\ndocumented composite, zoned diabase sills occurring in the outer zone, at Luodian in\nsouthern Guizhou63. Overall, the prevailing view on the distribution of intrusions in the\nEmeishan LIP is that they are mainly concentrated in the inner zone.\nThe timing of the Emeishan LIP intrusions broadly overlaps the main basaltic eruptions\nand extends from the earliest magmatic records to the waning stage (Supplementary\nFig. 3d; refs. 14-18,52-54,61,63,64). Although the current consensus is that Emeishan LIP\nextrusive and intrusive activities were essentially coeval, the analytical quality and\nprecision of available age determinations for the intrusions are insufficient to establish\na precise intrusive chronology comparable to that of the Siberian Traps.\nThe country rocks intruded by the Emeishan bodies are predominantly carbonates. In\nthe inner zone, intrusions were emplaced largely into Neoproterozoic (Sinian) dolostones that are nearly pure (containing very little clay or quartz), interbedded with\nsiliceous limestones, marlstones and minor shales of the Dengying Formation65. In the\nouter zone, the Luodian sills intrude Early-Middle Permian Sidazhai Formation\ncarbonates63. The development of contact aureoles within these carbonate successions\nsuggests metamorphic decarbonation and localized carbonate partial melting of the\ncountry rocks during emplacement65,66.”\nSupplementary Information, Supplementary Fig. 3: “\nSupplementary Fig. 3 | Geological map of the Emeishan large igneous province and\ncorrelation of the Maokou Formation, modified from the ref. 10. (a) Isopachs of the\nMaokou Formation outline a subcircular domal structure that formed prior to the\nEmeishan LIP volcanism. Black dash-dotted lines indicate profiles across the Maokou\nFormation (A–I and K–R). Blue dashed lines represent the inner, intermediate, and\nouter zones, defined by the extent of erosion of the Maokou Formation. Faults are\nmarked as follows: F —Longmenshan thrust fault; F —Ailaoshan-Red River slip fault;\n1 2\nF -Xiaojiang fault; F —Xichang-Qiaojia fault. Profile A-I in (b) and profile K-R in (c)\n3 4\nrepresent a west-east and a south-north oriented traverse across the Emeishan LIP,\nrespectively. Each section is labeled with the thickness of the Maokou Formation. All\nsections maintain a consistent vertical scale, with the exception of section A. (d) Time\nframework of Emeishan LIP magmatism based on Zircon U-Pb dates2,11-20.”\nAnother potential issue is that the authors are considering the effects of the Emeishan\nLIP to be the primary cause of what may be more global changes. This may be the case,\nbut there are other potential candidates in the Permian for large-scale changes in the\ncarbon cycle. The Emeishan LIP occurs at the end of the Late Paleozoic Ice Age (LPIA),\nwhich recent studies have suggested may have been instigated and ended by the\navailability of land within the tropical high-weatherability zones as Pangea moved\nslowly northwards (e.g., Goddéris et al., 2017; Marcilly et al., 2025). I think it is\nimportant to frame the carbon cycle instabilities across the Capitanian-Wuchiapingian\nin a broader context of widespread carbon fluctuations across the LPIA in the\nCarboniferous and Permian. Moreover, the placement of Emeishan within the high\nweatherability zones is important for the potential rate and extent of weatherability for\na domed sequence. Including this perspective could enhance the author’s findings of\nrapid unroofing above a mantle plume.\nWe thank the reviewer for this insightful comment and fully agree that placing the\nEmeishan LIP within the broader Late Paleozoic Ice Age (LPIA) weatherability\nframework helps refine our interpretation of the CO mechanisms. Accordingly, in the\nPaleozoic high-weatherability zones (HWZ), and (ii) explains how the Emeishan\nplume-driven uplift amplified this pre-existing weatherability. This framework\ndemonstrates that the Emeishan signal represents a plume-amplified weathering pulse\noperating within—and amplifying—broader LPIA weatherability patterns, rather than\nan isolated regional phenomenon. Our revised text now reads:\nLines 173–185:\n“In addition, the distinctive CO record of the Emeishan LIP is closely linked to its\n2\nbroader Late Paleozoic weatherability context. Marcilly et al. (ref. 41) reconstruct high-\nweatherability zones (HWZ) from climatically sensitive lithologies and highlight their\nimportance for the evolution of the Late Paleozoic Ice Age (LPIA). During the Permian,\nHWZ preferentially shifted toward the China blocks and Siberia rather than the arid\ncentral Pangean interior41, placing South China within a warm, high-runoff belt\nconducive to strong CO uptake. In such low-latitude settings, tectonic uplift can shift\n2\nweathering from supply-limited to kinetically controlled regimes and thereby\nenhancing CO consumption42. Within this framework, plume-related doming above\nplatform, producing a pronounced CO drawdown. Therefore, our Emeishan record\n2\nreflects a focused, characteristic, plume-amplified weathering pulse operating within\nglobal LPIA-scale changes in weatherability”\nThe calculations on the rate and volume of carbonate weathering from uplifted strata\nseem a little simplistic. The combined surface reservoirs of carbon in the Permian were\nmuch bigger than today, and the negative feedbacks that counteracted pCO2 increase\nwere different as there wasn’t a full-scale pelagic carbon pump set up yet (e.g., Ridgwell\nand Zeebe, 2005). A little more discussion is needed in how the unique forcings\nassociated with Emeishan (location, uplift amount, size) interplayed with the late\nPaleozoic carbon cycle in particular. This directly feeds into the catchy title of “Every\nLIP is different”, and some more thorough investigation of why each LIP is different\nwould be a welcome addition to the paper.\nWe thank the reviewer for the valuable comments. Overall, the reviewer’s main\nconcerns are (i) the simplification in our weathering calculation, (ii) the lack of\nPaleozoic carbon-cycle context (e.g., absence of a fully developed pelagic pump),\nand\n(iii) the need for additional discussion on how the specific forcings of the Emeishan\nLIP (e.g., location, size) relate to the paper’s central theme that “Every LIP is different”.\n(i) We have now included a sensitivity analysis in the Methods and added a statement\nacknowledging the limitations of our calculation.\nLines 630–639:\n“Sensitivity analysis of net CO consumption. To assess robustness, we quantified\n2\nhow the carbonate-weathering sink scales with the geometry of the uplifted dome and\nthe organic-carbon content by exploring dome radius of 600–900 km, average erosion\nthickness of 50–300 m, and bulk TOC values of 0.5–2 wt%. For all combinations, the\nnet carbon sink ranges from 1.41 × 104 Gt C at (600 km, 50 m, 2%) to 2.19 × 105 Gt C\nat (900 km, 300 m, 0.5%), and is therefore 103 Gt C under all tested conditions\n(Supplementary Fig. 6). Assuming a 2–4 Myr weathering interval, mean removal rates\n≫\nspan 0.003–0.11 Gt C yr-1. Note that our estimate of the carbonate-weathering volume\nand rate represents a simplified, first-order approximation intended to illustrate the\nlikely range of values rather than capture the full complexity of the system.”\nSupplementary Information, Supplementary Fig. 6: “\nSupplementary Fig. 6 | Sensitivity of the integrated net CO\n2\nconsumption potential\nfrom uplift-enhanced weathering of exposed Maokou carbonates, shown as a function\nof dome radius (km), average erosion thickness (m), and bulk TOC (wt%). Color scale\nindicates changes in TOC. The minimum across the parameter grid is indicated by a\nthicker yellow cross.”\n(ii) Following the reviewer’s suggestion, we now place the Emeishan weathering pulse\nwithin the Late Paleozoic (Neritan) carbon-cycle framework and have added a\ndiscussion noting that a pelagic deep‑sea CaCO sink had not yet fully developed before\n3\nthe Mid‑Mesozoic. Consequently, carbonate‑compensation feedback was weak in our\ninterval. Hence, uplift‑enhanced carbonate weathering would have exerted a larger and\nlonger‑lived downward pressure on atmospheric pCO than in the modern system,\n2\namplifying the impact of Emeishan domal erosion. Our revised text now reads (Lines\n196–208):\n“Notably, the Late Permian predates the Mid‑Mesozoic proliferation of pelagic\ncalcifiers and the establishment of a well-buffered deep‑sea carbonate sink. In such a\n‘Neritan’ ocean state, carbonate burial was dominated by platforms and shelves, and\nthe negative feedback associated with deep-sea carbonate compensation was inherently\nweaker and slower44. Consequently, a large, transient alkalinity–DIC injection from the\nrapid exposure and weathering of the Maokou carbonates would not have been\neffectively buffered by pelagic dissolution or precipitation. This inefficiency likely\nproduced a stronger atmospheric pCO response than in the modern ocean, sustaining\n2\npCO drawdown on Myr timescales and amplifying the climatic consequences of\n2\nenhanced carbonate weathering during the crustal doming of the Emeishan LIP. This\nPermian carbon-cycle architecture strengthens the plausibility of our interpretation that\ntectonic uplift may have dominated the atmospheric signal across the Emeishan LIP\ninterval.”\n(iii) We have now added a conclusion that synthesizes the distinctive features of the\nEmeishan LIP to better highlight its uniqueness. The revised text now reads (Lines 307–\n316):\n“More broadly, the climatic expression of LIP volcanism is shaped by a combination\nof factors, including emplacement style (extrusive vs. intrusive), mantle volatile\nloading, lithospheric structure, paleogeography, and the prevailing ocean–carbon state,\netc. The Emeishan LIP developed in a tropical, high-weatherability setting, where\ncrustal uplift exposed kilometer-thick Yangtze platform carbonates to intense erosion\nand weathering under a Late Paleozoic ocean characterized by weak pelagic buffering,\nwhile the associated basaltic magmatism was notably CO -poor. Together, these factors\n2\nexplain why the Emeishan LIP is distinct among continental flood basalts in exhibiting\npCO decline during the main basaltic phase.”\nLine 69: Change “biomarker” to “biomarkers”\nThank you. Revised accordingly.\nLine 182: Extensive uplift prior to the emplacement of a LIP is not a phenomenon\nunique to Emeishan. In particular, the North Atlantic Igneous Province (NAIP) has\nsubstantial uplift preceding the main phase of emplacement (Maclennan & Jones, 2006;\nConway-Jones & White 2022). The radius and uplift area for Emeishan are\nconsiderably less than calculated for the NAIP, which may be indirect evidence of a\nsmaller LIP volume. Therefore, the location of the uplift (i.e., whether it is in a high\nweatherability zone) is going to be critical for supporting the conclusion that rapid\ncarbonate weathering led to intense carbon release.\nWe thank the reviewer for pointing this out. We have now explicitly incorporated the\nhigh-weatherability zone (HWZ) mechanism into the discussion (Lines 173–185; see\nalso our response to the second main comment) and have also summarized this key\nfactor in the conclusion (Lines 307–316; see also our response to the third main\ncomment), where we synthesize the distinctive features of the Emeishan LIP.\nLine 234: A value of 135 ppm is way below CO2 concentrations in all modern systems.\nThis suggests a preservation issue rather than a true representation of magmatic CO2\ndegassing. As mentioned in the main point, it is thermogenic degassing that is likely to\nbe the most important source.\nWe agree that the value of 135 ppm reported by Tang et al. (2021) is relatively low and\nmay have been affected by preservation. We now explicitly state that this value should\nbe treated as a lower bound rather than a direct indicator of primary magmatic degassing.\nWe use it as a qualitative indicator consistent with our inference that Emeishan CO\n2\ninputs were not dominantly magmatic. As the reviewer also emphasized in the first\ncomment, thermogenic degassing has already been added to the main text.\nThe clarifying statement has now been added to Lines 232–237:\n“Note that the value of 135 ppm is lower than typical values observed in modern\nsystems and may have been affected by loss during preservation. Therefore, this value\nshould be interpreted with caution as a direct quantitative measure of magmatic\ndegassing. Nevertheless, it provides strong comparative evidence that the Emeishan\nbasaltic magmatism was notably low in CO and that degassing was not primarily\n2\nmagmatic.”\nFigure 3: The schematic drawing of the LIP has a stratovolcano in the centre of the LIP,\nwhich is not accurate. I suggest redrawing to show Hawaiian/Icelandic type eruptions\nfrom fissures or a shield volcano.\nFollowing the reviewer’s suggestion, we have now redrawn the LIP in Figure 3 and\nadded a clarification to the figure caption, indicating that this figure is intended as a\nconceptual illustration and does not reproduce the full geometry or complexity of the\nLIP. Additionally, the panel order and font size in the figure have been adjusted as\nsuggested by Reviewer #2.\nOur revised Figure 3 and its caption now read (Lines 952–964): “\nFigure 3 | Time series and evolutionary stages of Emeishan LIP emplacement, and\nthe carbon cycle spanning from the Guadalupian to the Lopingian.\n(a) Stable carbonate platform prior to the mantle plume formation. (b) Kilometer-scale\ncrustal uplift and erosion due to the Emeishan mantle plume during the Stage I, with\nthe orange dashed area representing the weathered Maokou Formation, leading to the\ndecrease of CO . (c) The main eruption of the Emeishan flood basalts during the Stage\n2 2\neruptions and associated greenhouse gas emissions (e.g., CO ) during the Stage III. The\n2\nwidths of the curved arrows with isotopic signals in (b, c) represent the magnitude of\ncarbon flux. Note that the thicknesses of Earth’s internal layers in this figure are not to\nscale. This figure is intended as a conceptual illustration and does not reproduce the full\ngeometry or complexity of the LIP.”\nLine 296: This statement needs revisiting through the lens of when and how much of\nthe Emeishan LIP was intrusive, and whether there was significant thermogenic\ndegassing associated with it.\nWe appreciate the reviewer’s comment. Please refer to our response to the first major\ncomment above and the corresponding revisions in the manuscript. In addition, this\nparagraph has now been revised accordingly, incorporating the reviewer’s three main\ncomments to refine our explanation of the Emeishan CO record.\nKey Message: This paper used a phytane-based phytoplankton proxy to reconstruct\npCO2 values across the Emishan LIP event during the Permian. The authors found that\npCO2 varied during different stages of the LIP, rather than follow a general increase, as\nis the expected norm during a LIP. Specifically, pCO2 decreased during the initial and\nmain stage of the LIP and only increased in the final stage. Significant increases in\nchemical weathering, associated with uplift, is suggested as the cause of the initial\ndecrease.\nSignificance: This paper provides detailed d13C_carb, d13C_org, and d13C_phytane\nmeasurements, the latter of which have a very limited published record in the Permian.\nFilling in these gaps has enabled the authors to reconstruct pCO2 over this time period\nusing the phytoplankton pCO2 proxy based on phytane. While this would be significant\non its own, the authors also discovered an unexpected decrease in the pCO2 values\nacross the Emishan LIP, providing nuance to our understanding of how LIPS influenced\nthe carbon cycle.\nWe thank the reviewer for the positive comment. Below, we addressed the reviewer’s\nconcerns point by point.\nData/Methods: The authors used d13C_carb, d13C_org, and d13C_phytane to calculate\nthe amount of carbon fractionation occurring in phytoplankton (epsilon_p). This value\nis then used to calculate pCO2 using previously established methods as mentioned in\nthe paper. The data and methods are sound, however there are some potential gaps in\ndetermining the quality of the data, specifically:\n1) Given the very low TOC values of many of the samples, how well was phytane\nisolated from the background signal?\n2) Was there any evidence of terrestrial input in the samples that may have altered the\norganic content?\nIn addition, Graham et al. (2025) recently calibrated epsilon_p of phytol to pCO2 in\nmodern sediments and applied their calibration to phanerozoic data sets, including\nphytane data sets. The authors should consider utilizing this calibration or address why\nthey have chosen not to use it.\n1) We thank the reviewer for this important point. To maximize biomarker yield, we\nused >150 g of powdered rock per sample for lipid extraction and performed three ASE\nextraction cycles until the extract color became lighter. Compared to typical geological\nsamples that require ~10–20 g, our starting mass was >10× larger, ensuring sufficient\nphytane abundance for duplicate isotopic analyses. To achieve baseline separation of\nphytane, each sample was treated with 5 Å molecular sieves to remove straight-chain\nn-alkanes. Before isotopic analysis, we verified phytane identity and purity by GC-MS\nand required baseline-resolved phytane peaks with good peak shape. Any\nchromatogram showing peak overlap (coelution) was excluded. All reported δ13C\nphy\nvalues are based on duplicate GC-IRMS injections to ensure data quality. Overall, the\nlarge starting mass, repeated extractions, n-alkane removal, explicit peak-purity criteria,\nand replicate measurements demonstrate robust isolation of phytane from background\nin our samples. In the revision, we now have added the above information to the\nMethods and included a representative GC-MS chromatogram illustrating phytane\nidentification and baseline separation (Supplementary Fig. 7).\nOur revised text in the Method now reads:\nLines 394–399: “Samples were extracted using a Thermo Scientific™ Dionex™\nASE™ 350 Accelerated Solvent Extractor in 9:1 dichloromethane/methanol mixture at\n100 °C to obtain total lipid extracts (TLEs). To maximize biomarker yield, more than\n150 g of powdered rock was used per sample for extraction, and three ASE extraction\ncycles were performed until the extract color became lighter, ensuring sufficient\nbiomarker abundance for subsequent isotopic analyses.”\nLines 402–408: “To remove n-alkanes, these fractions were transferred to cyclohexane\nand treated with activated 5Ǻ molecular sieves at 85 °C for 24 hours. Before isotopic\nanalysis, we verified phytane identity and purity by GC-MS and required baseline-\nresolved phytane peaks with good peak shape. Any chromatogram showing peak\noverlap (coelution) was excluded. Representative GC-MS chromatograms illustrating\nphytane identification and baseline separation are shown in Supplementary Fig. 7.”\nSupplementary Information, Supplementary Fig. 7: “\nSupplementary Fig. 7 | Gas chromatograms of representative Shangsi samples across\nthe Guadalupian-Lopingian boundary showing phytane identification and baseline\nseparation.”\n2) To address this important concern, we performed additional biomarker analyses on\nall samples to independently assess terrestrial organic matter input. Specifically, we\nmeasured n‑alkane distributions and calculated three established indices: the\nterrigenous-aquatic ratio (TAR), carbon preference index (CPI), and pristane/phytane\nratio (Pr/Ph). These analyses expand the geochemical dataset beyond the original\nphytane δ13C measurements. For the n-alkane analysis, all F1 fractions had been used\nfor molecular sieve separation to obtain sufficient phytane. Therefore, during the\nrevision, we re-extracted samples from powdered rock, performed column\nchromatography, and analyzed the n-alkanes by GC-FID. The results show that the\norganic matter in the Shangsi samples mainly reflects marine sources.\nThe TAR contrasts long-chain n-alkanes (C + C + C ; higher-plant leaf wax) with\n27 29 31\nshort-chain n-alkanes (C + C + C ; algae/bacteria). TAR > 1 indicates enhanced\n15 17 19\nterrestrial input, whereas TAR < 1 indicates aquatic/marine dominance (Meyers 1997).\nOur TAR results show that 33 out of 39 samples have values below 1, with an average\nof 0.58 for these samples. The six samples with TAR values greater than 1 are not\nconcentrated in any specific stratigraphic level but occur randomly throughout the\nsection. The CPI quantifies odd-over-even predominance in long-chain n-alkanes.\nFresh higher-plant leaf waxes show strong odd-over-even predominance (CPI 1),\nwhile substantial algal inputs or thermal maturity yield values around 1 (Bray and Evans\n≫\n1961; Peters et al. 2005). Our CPI values range from 0.67 to 0.98 (average = 0.89 ±\n0.07, n = 39), approaching 1, indicating no obvious odd-over-even dominance\ncharacteristic of fresh higher-plant wax. In addition, Pr/Ph values < 1 are generally\ninterpreted as typical of anoxic/dysoxic marine settings, values ≈ 1–3 are common in\nreducing but non-sulfidic siliciclastic shelves, and values > 3 characterize\nfluvial/deltaic sediments with high terrestrial input (Hughes et al., 1995). In our samples,\nPr/Ph ranges from 0.23 to 1.03 (average = 0.59 ± 0.14; n = 38), with no values > 1.5\nand none approaching > 3. This distribution is consistent with reducing marine\ndeposition. The Pr/Ph data are part of another manuscript focusing on redox conditions,\nwhich is currently under review and therefore remains unpublished.\nWe thank the reviewer for this comment. To clarify this point and address the reviewer’s\nconcern, we have added the new data and above discussion to the revised manuscript.\nLines 420–439: “Terrestrial organic matter input. To assess the contribution of\nterrigenous organic matter in the Shangsi samples, we used n‑alkane and isoprenoid\nindices (Supplementary Table 2). The abundances of these compounds were determined\nusing an Agilent 8890 gas chromatograph equipped with a flame ionization detector\n(GC-FID) and a DB-5 capillary column (30 m × 0.25 mm × 0.25 μm). The temperature\nprogram started at 70 °C, then increased at a rate of 10 °C/min to 210 °C, followed by\n3 °C/min to 300 °C, where it was held for 30 min. Based on peak areas, we calculated\n(i) the terrigenous-aquatic ratio (TAR; TAR = [C + C + C ]/[C + C + C ]),\n27 29 31 15 17 19\nwhich contrasts long‑chain (C –C ) higher-plant wax alkanes with short‑chain (C –\n27 31 15\nC ) algal/bacterial alkanes68, (ii) the carbon preference index (CPI; CPI = [C + C +\n19 25 27\nC + C ]/[C + C + C + C ]) of long‑chain n‑alkanes as an indicator of\n29 31 24 26 28 30\nodd‑over‑even predominance typical of fresh higher‑plant waxes69,70, and (iii)\npristane/phytane (Pr/Ph), which generally reflects redox and depositional setting71. At\nShangsi, 33 out of 39 samples have TAR values below 1, with an average of 0.58 for\nthese samples. The six samples with TAR values greater than 1 are not concentrated at\nany particular stratigraphic level but occur randomly throughout the section.\nLong‑chain CPI values cluster around 1 (0.67–0.98; average = 0.89 ± 0.07, n = 39), and\nPr/Ph ranges from 0.23–1.03 (average = 0.59 ± 0.14; n = 38), with no values exceeding\n1.5. Together, these indices indicate that the organic matter in the Shangsi samples is\nprimarily derived from aquatic/marine sources, with a relatively minor higher-plant\ncontribution.”\n3) We thank the reviewer for highlighting Graham et al. (2025), which provides a\nmodern‑ocean calibration of the phytol ε -[CO (aq)] relationship (Deming regression;\np 2\nε ≈ 23.9‰, b ≈ 82.4) and has encouraged the wider use of phytol/phytane‑based pCO\nf 2\nreconstructions. We now explicitly acknowledge this contribution in our Methods and\nclarify how it relates to our parameterization.\nOur study parameterizes b from phosphate following established practice (b = 82.05 ×\n[PO 3−] + 86.28), which for [PO 3−] = 0.2 µM gives b ≈ 102. This value is well below\n4 4\nthe legacy constant b = 170 and in the same “low‑b” range as modern calibrations\n(Graham et al., 2025). As stated in our Methods, reasonable choices of b mainly shift\nabsolute pCO levels and do not alter the temporal pattern we reconstruct for the\nRegarding the choice of ε, we respectfully retain ε = 28‰ for the following reasons:\nf f\n(i) In our dataset, the εₚ values are very close to 23.9‰. Because the εₚ-pCO function\n2\nis asymptotic, setting ε at 23.9‰ would place more data near the ε ≈ ε limit, where\nf p f\nsmall uncertainties in ε propagate into disproportionately large or indeterminate pCO\np 2\nvalues. Graham et al. (2025) state that the reason for the low ε is the lack of very high\nf\nCO (aq) samples in the modern calibration set. The refined Deming calibration\n2\nexcludes high-CO (aq) coastal sites (to avoid terrigenous organic matter), so the dataset\n2\nlacks the high-CO (aq) end needed to achieve maximum fractionation, thereby\n2\nlowering the fitted intercept representing ε.\nf\n(ii) Widespread expression of CCMs and active HCO - uptake (β‑carboxylation) in\n3\nmodern phytoplankton communities depresses apparent fractionation, again tending to\nlower the empirical ε under present‑day conditions. Graham et al. (2025) suggest that\nf\nε may have been higher during deep-time intervals with CCMs less active, implying\nf\nthat 23.9‰ is not necessarily a universal constant across all time or taxa. The Permian\nrepresents a transition from the Late Paleozoic icehouse to a hothouse climate, with the\nmid-Permian generally warmer and the late Permian hotter. Under higher CO levels,\nf\nhigher, as also discussed by Graham et al. (2025).\nWe appreciate the reviewer pointing this out, and we have now incorporated the above\ndiscussion into our Methods. Our revised text now reads:\nLines 505–525: “A recent modern‑ocean calibration of the phytol ε -[CO (aq)]\np 2\nrelationship, based on a Deming regression, reports ε ≈ 23.9‰ and b ≈ 82.4 (ref. 24).\nf\nThis study refines and strengthens the phytol/phytane‑based approach to pCO\n2\nreconstruction, thereby improving its reliability and expanding its applicability for\nfuture studies. In our study, the b value is parameterized from the phosphate relationship,\nwhich yields b ≈ 102. This value is substantially lower than the traditional constant of\nb = 170 and broadly consistent with the low‑b range in modern calibrations24. As for ε,\nf\nwhile the modern calibration provides an important reference framework, several\nconsiderations suggest that its ε value may not be directly applicable to our study. First,\nf\nthe calibration domain of modern datasets is limited at the high‑[CO (aq)] end. The\n2\nopen ocean rarely attains very high [CO (aq)], and near-shore waters, where [CO (aq)]\n2 2\ncan be higher, are typically excluded to minimize terrestrial inputs. This restriction\nreduces the opportunity to observe maximum fractionation, thereby leading to a lower\nfitted ε (ref. 24). Second, carbon‑concentrating mechanisms (CCMs) and active HCO -\nf 3\nuptake (β ‑carboxylation) are pervasive in modern phytoplankton and tend to depress\nthe apparent fractionation. During deep-time intervals, these pathways were likely\ndown‑regulated under elevated pCO\n2\n, implying a higher effective ε\nf\n. In addition,\nbecause of the asymptotic nature of the ε -pCO relationship, adopting a low ε places\np 2 f\nmore of our data near the ε ≈ ε limit, where small uncertainties in ε can propagate\np f p\ninto disproportionately large pCO (ref. 24). For these mechanistic and methodological\n2\nreasons, we adopt ε = 28‰ in our pCO reconstructions.”\nf 2\nClarity/context: The paper is well written and delivers its key messages clearly. Figs 1\nand 2 are easy to interpret and do a good job of condensing and comparing multiple\nlines of evidence. The small volcano in the bottom left corner of fig 2 is a bit\nconfusing and doesn’t match with the rest of the fig layout.\nWe now have removed the small volcano symbol to for better overall consistency of the\nfigure.\nFIG 3 - The order of fig 3 is also a little confusing as the bottom section is labeled ‘a’\nand the top ‘d’. Most readers will read the fig from top to bottom, and the labels should\nbe switched to reflect this order, or both the figure and labels should be switched (i.e.\ntop = ‘a’ pre-mantle; bottom = ‘d’ stage III). A timeline attached to the side might also\nclarify the order in which the figures appear, so that the audience immediately knows\nwhich are the older stages and which are the younger.\nThe text in this figure is also very small and it is unclear why some text is distinguished\nby colored boxes behind it and some is not.\nIt is also not clear why Stage III specifies composite weathering input ‘dominant crust-\nderived carbon’ while Stage I and II separate organic and carbonate weathering inputs.\nThe relevance of the thickness of the carbon arrows should be noted in the figure, not\njust the description.\nThe dashed line mentioned in the text description is orange, not red.\nWe thank the reviewer for these helpful suggestions. All requested changes have been\nmade in the revised Fig. 3.\n(i) The panel order has been re-arranged so that figure now reads a→d from top to\nbottom.\n(ii) All text in Fig. 3 has been enlarged to improve clarity and legibility. Some of the\ntext is highlighted with colored boxes because the figure contains many annotations.\nThese colors are used simply to distinguish the text describing the geological processes\nand characteristics of each stage.\n(iii) The flux arrow in panel d has been removed because this is not discussed in the\nmanuscript. For panels b and c, we have added labels directly next to the arrows so that\nthe thick arrows are marked as “high flux” and the thin arrows as “low flux”, making\nthe meaning of arrow thickness explicit.\n(iv) The dashed-line color in the caption of panel b has been corrected to orange.\n(v) Additionally, the stratovolcano in the LIP schematic has been replaced with\nHawaiian/Icelandic-type eruptions from fissures or a shield volcano, as suggested by\nReviewer #1.\nOur revised Figure 3 and its caption now read (Lines 952–964): “\nFigure 3 | Time series and evolutionary stages of Emeishan LIP emplacement, and\nthe carbon cycle spanning from the Guadalupian to the Lopingian.\n(a) Stable carbonate platform prior to the mantle plume formation. (b) Kilometer-scale\ncrustal uplift and erosion due to the Emeishan mantle plume during the Stage I, with\nthe orange dashed area representing the weathered Maokou Formation, leading to the\ndecrease of CO . (c) The main eruption of the Emeishan flood basalts during the Stage\n2 2\neruptions and associated greenhouse gas emissions (e.g., CO ) during the Stage III. The\n2\nwidths of the curved arrows with isotopic signals in (b, c) represent the magnitude of\ncarbon flux. Note that the thicknesses of Earth’s internal layers in this figure are not to\nscale. This figure is intended as a conceptual illustration and does not reproduce the full\ngeometry or complexity of the LIP.”\nExpertise: I have kept most of my comments related to the organic geochemistry and\ncarbon-isotopes section of this paper as that falls within my expertise. I do not have\nstrong expertise in Permian LIPs and mantle dynamics and as such have a limited\nassessment of those sections.\nWe sincerely thank the reviewer for the constructive comments and suggestions, which\nhave greatly improved our manuscript."}, {"round": 2, "reviewer_comments": {"reviewer_2": "The authors have addressed my main concerns from the initial review through the following:\n1. Including chromatograms showing that phytane was clearly separated from the background signal\n2. Adding CPI data from new analysis showing low inputs of terrestrial contamination\n3. Expanding on their methodology for extractions\n4. Expanding on their justification for their choice of epsilon values\nThese additions are both comprehensive and satisfactory. The authors have greatly expanded on their methodology and\nadded new data which alleviate doubts about the quality of the phytane measurements. The paper is a well written argument\nfor a newly discovered relationship between the Emishan LIP and the Permian carbon cycle. I do not have any further\ncritiques."}, "author_response": ""}]

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