Rock weathering can counteract river CO2 emissions induced by permafrost thaw

Abstract

Climate-induced permafrost thaw unlocks large stores of organic carbon that are mineralized and emitted as carbon dioxide (CO₂) from rivers to the atmosphere¹. Concurrently, warming and permafrost thaw can increase mineral weathering rates, thus affecting the release and sequestration of inorganic carbon^2,3,4. Yet how these biological and geological carbon cycles interact and jointly affect CO₂ dynamics (emission compared with drawdown) in permafrost rivers remains unknown⁵. Here we combine CO₂ emissions, organic and inorganic solute concentrations, dual carbon isotopes (δ¹³C–Δ¹⁴C) and geochemical modelling to infer how permafrost thaw may affect river biogeochemistry over decades to centuries across the Qinghai–Tibet Plateau. Leveraging a gradient of thermal permafrost degradation, we find that river CO₂ emissions decline, whereas solute fluxes from rock weathering increase with decreasing permafrost cover. Across this region, net CO₂ drawdown fluxes from rock weathering are about 35% of river CO₂ emissions, varying from around 15% in catchments with continuous permafrost to more than 100% in catchments with discontinuous or isolated permafrost. Thus, carbon fluxes from chemical weathering may become increasingly important with ongoing permafrost thaw, potentially even outpacing river CO₂ emissions. Our findings disentangle the interplay between biological and geological carbon fluxes that are important for the cryosphere and the global carbon cycle.

Main

River networks perfuse and integrate landscapes⁶, receiving organic carbon (OC) from terrestrial primary production and dissolved inorganic carbon (DIC) \(({\rm{D}}{\rm{I}}{\rm{C}}=\sum {{\rm{C}}{\rm{O}}}_{2({\rm{a}}{\rm{q}})}+{{{\rm{HCO}}}_{3}}^{-}+{{{\rm{CO}}}_{3}}^{2-})\) from soil respiration and chemical rock weathering. This OC is metabolized in rivers and returned to the atmosphere as carbon dioxide (CO₂) or methane (CH₄), buried in sediments, or transported to coastal oceans. Similarly, DIC can precipitate, outgas or travel downstream^7,8. Global carbon inputs to rivers (about 3.2 petagrams of carbon (PgC) yr⁻¹), the resulting riverine carbon emissions (around 2.3 PgC yr⁻¹) and carbon export to the ocean (approximately 0.8 PgC yr⁻¹)⁸ are on par with the net terrestrial carbon sink (that is, the net primary productivity minus respiration without considering land-use changes; about 3.2 PgC yr⁻¹)⁹. Consequently, rivers forge important connections between terrestrial, marine and atmospheric carbon pools that modulate the global carbon cycle⁸. Despite major advances in understanding the role of rivers in the carbon cycle, key knowledge gaps persist regarding the interaction between biological and geological carbon sources and sinks⁵, as well as their responses to climate change.

Of primary concern is the impact of climate warming on the vast OC stocks (approximately 1,014 PgC) stored in the Northern Hemisphere permafrost soils¹⁰. Permafrost thaw is mobilizing a portion of this large, near-surface OC reservoir on Earth¹¹. The subsequent breakdown of this ancient, biolabile OC in rivers fuels in situ respiration and methanogenesis, thus amplifying emissions of aged CO₂ and CH₄, and potentially perturbing the contemporary carbon cycle^{12,13,14,15,16,17,18}. However, permafrost thaw has effects beyond mobilizing ancient OC. In particular, recent studies have demonstrated that multi-decadal increases in DIC fluxes and associated dissolved solutes in permafrost rivers could reflect intensified chemical weathering^2,3,4,19,20. Given that landscape-scale weathering fluxes are controlled by a combination of lithology²¹, hydrology^22,23,24, temperature^25,26 and exposure of reactive surfaces^27,28, climate warming and permafrost thaw could have multifaceted impacts on these fluxes^2,3,4. Whereas climate warming can directly increase dissolution kinetics²⁶, the associated permafrost thaw may alter weathering reactions by exposing minerals or by modulating the hydrology and promoting acidic redox conditions in sub-surface environments⁴. Importantly, weathering reactions in the critical zone affect biogeochemistry through the hydrosphere because groundwater transports weathering products—including carbon—to fluvial ecosystems.

The role of weathering reactions for the carbon cycle depends on the nature of the exposed minerals^21,29. Weathering of silicate and carbonate minerals by carbonic acid produces alkalinity and draws down CO₂ (refs. ^30,31). By contrast, weathering of sulfide minerals—such as pyrite—produces sulfuric acid and releases CO₂, either by interacting with the alkalinity pool or by dissolving carbonate minerals^32,33. In reality, biological and geological CO₂ fluxes are intertwined and co-vary across permafrost rivers. For example, acidity produced by respiration of OC could be buffered by alkalinity production from chemical weathering²⁸. These interactions may be particularly important in mountainous landscapes with widespread permafrost coverage that are strongly exposed to climate change and characterized by rapid weathering rates^34,35. Efforts to isolate the impact of permafrost thaw on chemical weathering are rare, and no studies have quantified the balance between biological and geological CO₂ sources and sinks in rivers draining permafrost landscapes subject to climate warming. These knowledge gaps introduce substantial uncertainties in understanding the contemporary and future role of permafrost landscapes and their rivers within the global carbon cycle⁸.

Here, we examine the balance of riverine CO₂ fluxes resulting from OC mineralization and chemical weathering spanning a permafrost gradient on the Qinghai–Tibet Plateau (QTP; Fig. 1a). The QTP is the largest contiguous cryosphere outside the Arctic and Antarctic³⁶, with an estimated OC storage of 50.4 Pg in its permafrost soils³⁷, partly dating back to the late Pleistocene³⁸. The permafrost on the QTP is, on average, warmer and has a higher mean active-layer thickness (about 0.21 m) compared with Arctic permafrost (about 0.10 m), with 44% of permafrost at a mean ground temperature above −0.5 °C, which makes it particularly vulnerable to climate warming³⁹. From the cold plateau in the north and west of the study area, air temperatures increase eastward and southward, and permafrost extent decreases from continuous to discontinuous, sporadic and isolated (Fig. 1a and Extended Data Fig. 1a). By substituting space for time⁴⁰, this landscape and its rivers potentially hold information about the sensitivity of biogeochemical cycles to climate warming and permafrost thaw over decadal to centennial timescales, timescales of permafrost landscape changes that are impossible to observe with long-term monitoring approaches.

**Fig. 1: Environmental gradients on the QTP.**

A permafrost gradient on the QTP

We repeatedly sampled 50 rivers in the headwaters of the largest river systems of Asia (Yellow, Yangtze, Lancang, Nu, Derung, Yarlung Tsangpo, Marja Tsangpo and Indus Rivers; Fig. 1a) during the ice-free season between April and October over more than 3 years. The sampled headwaters cover a total area of approximately 780,000 km², span elevations from 1,650 m to 4,820 m above sea level, and drain all major lithologies of the QTP, including a variety of sedimentary, igneous and metamorphic rocks (Extended Data Fig. 1b and Supplementary Tables 1 and 2). We include basins on the cold plateau underlain by continuous permafrost and catchments draining warmer landscapes with discontinuous, sporadic or isolated permafrost cover.

Across these landscape-scale environmental gradients, we found that permafrost extent is inversely related to average annual air temperature (principal component 1, PC1; Fig. 1b,c and Extended Data Fig. 1a), consistent with thermal permafrost degradation in response to climate change⁴¹. Because permafrost extent tends to be the highest on the flat plateau in the northern and eastern parts of the study area, we also found that permafrost extent declines moderately with the catchment-averaged steepness index (k_sn), a proxy for landscape-scale erosion rates^42,43 (Fig. 1b,d and Extended Data Fig. 1c). The permafrost gradient does not substantially co-vary with factors that may affect chemical rock weathering (for example, lithology and average annual precipitation; Fig. 1b,e,f and Extended Data Fig. 1b,d). Furthermore, permafrost does not co-vary with wetland coverage, which can affect dissolved organic carbon (DOC) export to rivers⁴⁴ (Fig. 1b). Thus, our sampled catchments allow us to relate changes in permafrost extent and temperature to river biogeochemistry, mostly independently from other environmental variables that also affect weathering and carbon fluxes.

Solute and carbon sources in QTP rivers

To assess the concentrations and sources of carbon and inorganic constituents in QTP rivers, we measured CO₂ partial pressure (\({p}_{{\mathrm{CO}}_{2}}\)) and emission rates, concentrations of DOC, DIC, major dissolved cations and anions, and isotopes of carbon (δ¹³C–Δ¹⁴C), sulfur (δ³⁴S) and oxygen (δ¹⁸O). We found all 50 rivers to be supersaturated in CO₂ with respect to atmospheric equilibrium, with a median saturation of 230% (interquartile range 200–270%), and \({p}_{{\mathrm{CO}}_{2}}\) of 650 (580–750) μatm and a median riverine CO₂ emission rate of 95 (81–120) mmol m⁻² day⁻¹ (Extended Data Fig. 2). These emissions are bracketed by previous observations from QTP rivers (Supplementary Table 3), but remain below the median estimate (594 mmol m⁻² day⁻¹) for global rivers⁸. Concentrations of DOC and DIC are 380 (320–440) μmol l⁻¹ and 1,800 (1,400–2,600) μmol l⁻¹ (Extended Data Fig. 2 and Supplementary Table 2), respectively, whereas total dissolved ion concentrations (4,600 (2,900–5,900) μmol l⁻¹) are dominated by Ca²⁺ (750 (580–1,000) μmol l⁻¹), Mg²⁺ (480 (230–840) μmol l⁻¹), Na⁺ (410 (250–830) μmol l⁻¹), SO₄²⁻ (430 (260–730) μmol l⁻¹) and HCO₃⁻ (2,000 (1,300–2,600) μmol l⁻¹; Extended Data Fig. 2 and Supplementary Table 2). No substantial seasonal trends in either CO₂ or dissolved chemistry were apparent over the study period (Extended Data Fig. 3), although water temperatures increased from April to August and runoff increased slightly through summer and into fall (Extended Data Fig. 4).

We found that the dissolution of carbonates contributed the most cations (78% (70–86%) of the total cation charge), whereas silicates and evaporites contributed a smaller fraction (6.0% (3.6–9.5%) and 6.3% (1.6–14%), respectively; Extended Data Fig. 5a and Supplementary Table 2). With one exception, samples with more than 25% and up to 78% cation contribution by evaporites occurred exclusively in eight sites of the westernmost Yellow River, Yangtze River and Lancang River, and are probably controlled by evaporitic units within the metasediments⁴⁵. Coupled isotopic compositions of δ³⁴S-SO₄²⁻ and δ¹⁸O-(SO₄²⁻–H₂O) identify a mixing trend between riverine sulfate (SO₄²⁻) derived from pyrite and evaporite sources (Extended Data Fig. 5b). We infer that sulfate from sulfide oxidation balances 23% (11–36%) of the total cation charges (Extended Data Fig. 5a).

Interactions of organic and inorganic carbon pools

Based on dual carbon isotopes (δ¹³C–Δ¹⁴C) and a three-endmember mixing model⁴⁶, we found that old permafrost- and rock-derived carbon substantially contributes to DOC, DIC and CO₂ pools in QTP rivers. DOC is a mix of modern (67 ± 11%) and old permafrost (21 ± 13%) as well as rock (carbonate and petrogenic OC; 12 ± 6%) carbon sources, with an average age of 2,232 ¹⁴C years (Fig. 2a,b). DIC can be modelled as a mixture of modern (49 ± 19%) and old permafrost (34 ± 19%) and rock (19 ± 8%) sources, with an average age of 3,512 ¹⁴C years (Fig. 2a,b). Riverine CO₂ falls between DOC and DIC in both its δ¹³C and Δ¹⁴C signatures and has an average age of 1,655 ¹⁴C years, consistent with a mix of modern permafrost (66 ± 11%), old permafrost (22 ± 13%) and rock (12 ± 7%) carbon sources (Fig. 2a,b).

**Fig. 2: Carbon isotopes and unmixing.**