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Sustained and intensified lacustrine methane cycling during Early Permian climate warming - Nature Communications
Results and discussion .
Age and stratigraphic constraints .
Samples were collected from the Lucaogou Formation at the Jingjingzigou section along the southern margin of the Junggar Basin (Fig.? 1 and Supplementary Fig.? 1 ) for zircon U-Pb and geochemical analyses. The Lucaogou Formation can be subdivided into two members. The lower member mainly consists of mudstone, shale, dolomitic siltstone, and dolomite, with minor amounts of gypsum in some layers. The upper member is composed of organic-rich shale interbedded with dolomite beds and nodules without evaporite minerals (Fig.? 2a, c ; Supplementary Figs.? 1 and 2 ). This sequence reflects the evolution from a relatively shallow evaporative lake to a persistently deep brackish-to-freshwater lacustrine environment (see Supplementary Note? 1 ). Despite decades of sedimentological and geochemical/hydrocarbon research, due to the economic importance of the Lucaogou Formation 11 , 12 , 13 , 22 , 23 , the succession lacks any reliable age constraints in the absence of datable volcanic ash beds and biostratigraphically useful fossils 24 , 25 . Previous detrital zircon U-Pb geochronology obtained by in situ laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS) assigned broad maximum depositional ages of ca. 270–268 Ma 24 or ca. 261 Ma 26 to the Lucaogou Formation. However, limited accuracy due to reworked zircons and/or post-crystallization Pb loss can lead to statistically biased results. In addition, a previously published high-precision U-Pb CA-ID-TIMS age of 281.39?±?0.10 Ma 27 from the overlying Hongyanchi Formation from the southern Bogda Mountains (Figs.? 1 and 2c ) resulted in a contradictory stratigraphic framework.
Fig. 2: Permian stratigraphy and geochronology of the southern Junggar Basin. a Outcrop photograph of organic-rich shale interbedded with dolomite beds and nodules from the upper member of the Lucaogou Formation. The arrow points to the location of the volcanic ash bed (sample VA-1) sampled for zircon geochronology (inset shows a close-up view of the ash bed). b Concordia plot and 206 Pb/ 238 U ages of zircons analyzed using the U-Pb CA-ID-TIMS method; excluded analysis z4 shown in gray. Vertical bars represent 2σ analytical uncertainty of individual zircon analyses. c Stratigraphic column of the southern Junggar Basin (modified from ref. 11 ). Arrows indicate stratigraphic positions of dated ash beds (bentonites) and tuffaceous siltstone (blue–published ages of ref. 27 ; red–new ages presented in this study).
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Here we present high-precision U-Pb zircon age from a volcanic ash bed in the upper, organic-rich member of the Lucaogou Formation (Fig.? 2a, b ). The 4?cm-thick ash layer interbedded within shales occurs ~925?m above the base of the Lucaogou Formation (Figs.? 2 c and 3 ). The sample (VA-1) contains zircons that are small, equant or prismatic, and euhedral, with oscillatory zoning under cathodoluminescence (Supplementary Fig.? 3a ). The Th/U ratios of the zircon crystals vary from 0.26 to 1.27 (Supplementary Dataset? 1 ). The U-Pb ages determined by LA-ICP-MS have an average 2σ uncertainty of ±5.63 million years (Myr) and are distributed around a well-defined peak, with a weighted mean 206 Pb/ 238 U age of 286.14?±?0.65?Ma (2σ internal error only; mean-squared weighted deviation [MSWD]?=?1.01, n ?=?53; Supplementary Fig.? 3a ). For further verification, four single zircon grains from this sample were analyzed independently by the CA-ID-TIMS method (average 2σ uncertainty of ±0.55?Myr), with the three youngest analyses constituting a coherent cluster with a weighted mean 206 Pb/ 238 U age of 286.39?±?0.25/0.30/0.43?Ma (2σ; MSWD?=?2.0; Fig.? 2b, c and Supplementary Dataset? 2 ). Furthermore, one tuffaceous siltstone (sample TS-1) from the uppermost part of the underlying Jingjingzigou Formation was analyzed using the LA-ICP-MS method. Ninety-four zircon analyses from this sample yielded a wide range of ages, with a weighted mean 206 Pb/ 238 U age of 294.1?±?1.4?Ma (2σ; MSWD?=?1.5; Fig.? 2c and Supplementary Fig.? 3b ) based on 13 youngest analyses (YC2σ[3+] 28 ), and this is interpreted as being the maximum constraint on the depositional age. The available radioisotope geochronology collectively places the lower and upper boundaries of the Lucaogou Formation at ca. 294?Ma and ca. 285?Ma, respectively, and constrains its upper shale member to the Artinskian Stage; therefore, the age is significantly older than previous estimates 24 , 25 , 26 .
Fig. 3: Stable carbon isotopes, estimated land surface temperatures from the Junggar Basin, and comparisons with the Earth system changes during the Early Permian. The different colors in the stratigraphic column indicate changes in the lithology of the Lucaogou Formation. a Bulk organic matter δ 13 C org record. b Dolomite δ 13 C carbonate record. c C 29 and C 30 αβ hopane δ 13 C values. Error bars denote one standard deviation between duplicate analyses. d Chemical index of alteration (CIA) and land surface temperature (LST) estimates. The curves in ( a ) and ( d ) represent the seven-point moving averages. e CIA trend from the glacial to postglacial transition succession in the Karoo Basin of South Africa 21 with CA-ID-TIMS zircon age constraints 19 and temporal variations in low-Mg-calcite oxygen isotope (δ 18 O) values from low- and high-latitudinal fossil shells 50 , 51 . f Documented glacial deposits 20 and reconstructed global atmospheric partial pressure of CO 2 ( p CO 2 ) curve (75% confidence interval) 15 during the Early Permian. Timescale from an updated version (2022) of the International Chronostratigraphic Chart 72 .
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Organic carbon isotope excursions (CIEs) have been proven to be an effective global stratigraphic correlation proxy 29 . Our results show that a prominent negative δ 13 C org excursion occurs at the upper part of the succession, with a total CIE magnitude of ~3.5‰ (above ~725?m; Fig.? 3a and Supplementary Dataset? 3 ). The observed parallel δ 13 C TLE and δ 13 C Asph records from the total lipid extract (TLE) and asphaltene (Asph), after extraction and separation, also exhibit a largely similar negative CIE in shape and magnitude to that of the bulk δ 13 C org record. In addition, all the δ 13 C n -alkane records of short-chain n -C 19 , mid-chain n -C 21 , and long-chain n -C 27 alkanes (Supplementary Dataset? 4 ) display a negative CIE with a magnitude of ~4‰ (Supplementary Fig.? 4a ). It is considered unlikely that the thermal maturity (early oil window) and proportional changes in the organic matter resulted in the observed CIE in the Lucaogou shales (see Supplementary Note? 2 ; Supplementary Fig.? 5 ). Importantly, under our high-precision CA-ID-TIMS age constraint, the negative CIE is comparable to that recorded in coeval marine brachiopod shells (USA and Russia) 30 , carbonate (South China) 31 , and in coastal strata (North China; 29 Supplementary Fig.? 4 ). Therefore, the observed parallel CIE signatures in bulk δ 13 C org and δ 13 C n -alkane reflect a perturbation of the global carbon cycle during the Artinskian (see Supplementary Note? 2 ).
Intensified ecosystem-level microbial CH 4 cycling .
Investigating microbial CH 4 cycling in pre-Holocene environments is challenging owing to the scarcity of diagnostic proxy records; some lipid biomarkers (e.g., glycerol dialkyl glycerol tetraether and archaeol 32 , 33 ) may be invalid with an increase in the thermal maturity of the organic matter. In this study, we present the distinctive δ 13 C records of authigenic dolomites and hopanes (bacterial-derived biomarkers) from the Lucaogou Formation, which provide new insights into the metabolic activities of methanogens and methanotrophs in the lake ecosystem during the Early Permian.
Dolomite beds and nodules in the upper member (Fig.? 2a ) have very positive δ 13 C values (+5.8 to +16.0‰) that are significantly higher than the δ 13 C values (+5.3 to +8.3‰) of the dolomite in the lower member (Fig.? 3b and Supplementary Dataset? 5 ). Several mechanisms have been proposed for 13 C enrichment in inorganic C pools 34 , 35 , 36 . Of these, the Rayleigh distillation of volatile CO 2 under highly evaporative conditions 34 would not have been effective in paleo-Lake Junggar, as the upper member was not deposited in a hypersaline environment 12 , 13 . In addition, the photosynthetic fixation of CO 2 during productivity blooms cannot explain the positive values, because this process usually only enriches δ 13 C values by +2 to +3‰ 35 , and this Permian lake was not eutrophic 13 . Such positive δ 13 C signatures have been recently attributed to authigenic dolomite precipitation associated with microbial methanogenesis 10 , and it is likely that some dolomite samples with less positive δ 13 C and lower δ 18 O values in the upper member have been influenced by subsequent diagenesis 36 (Supplementary Fig.? 6 ). Microbial methanogenesis is geochemically characterized by significant C isotopic fractionation, generating 13 C-depleted biogenic CH 4 (δ 13 C as low as ?60 to ?110‰) 37 and 13 C-enriched CO 2 (δ 13 C up to +15‰ or higher) 36 . Such isotopically heavy CO 2 acted as a substantial C source and was incorporated into the authigenic dolomite. The closest modern analogues of these dolomites, commonly observed in organic carbon-rich continental margin sediments, have been documented in the Gulf of California 38 and along the Peru Margin 39 , where methanogenesis is highly active in oceanic sediments. Thus, the 13 C-enriched authigenic dolomites presented here are a fingerprint of biogenic CH 4 production in lake sediments.
Putative methanogen microfossils have been found in these 13 C-enriched dolomites from the adjacent Hongyanchi section, and their abundances show a positive correlation with δ 13 C values 10 . The elevated δ 13 C signature of the dolomites can therefore be used to trace changes in methanogenesis. In the current study, the dolomite δ 13 C values show an overall increasing trend from the bottom to the top within the succession (Fig.? 3b ). The dolomite beds and nodules that are marked by high δ 13 C values occur above ~610?m. Most importantly, the more abundant strongly 13 C-enriched dolomites occur within the upper part of the Lucaogou Formation (Fig.? 3b ), indicating a higher methanogenic rate and/or an expanded methanogen community in the anoxic lake sediments during this period. In contrast, the absence of exceptionally 13 C-depleted authigenic dolomite in the studied section (δ 13 C carbonate values typically? An earlier comprehensive study 13 revealed that the saturated hydrocarbon fraction from Lucaogou shales was depleted in 13 C, and this possibly indicates the presence of hopanes derived from methanotrophic bacteria. In this study, we conducted a compound-specific C isotope analysis of hopanes. Hopanoids are not exclusive to methanotrophs, but their stable C isotopic compositions can be used to assess specific methanotroph contributions 33 , 41 , 42 , 43 , 44 , 45 . Methanotrophic bacteria use biogenic CH 4 as a carbon source for the biosynthesis of membrane lipids (e.g., hopanoids) that are highly 13 C-depleted. Our results show that the hopanoids in Lucaogou shales are dominated by C 30 17α,21β-hopane and C 29 17α,21β-norhopane (Supplementary Fig.? 7 and Supplementary Note? 1 ). The hopane δ 13 C values remain low in all samples analyzed, ranging from ?44.1 to ?62.6‰ for C 30 17α,21β-hopane and ?41.6 to ?53.6‰ for C 29 17α,21β-hopane (Fig.? 3c ). The δ 13 C values in these two compounds yield a positive correlation (Supplementary Fig.? 5d ), indicating that they have a similar bacterial community source. Their corresponding 17β,21α(H) isomers are also characterized by similar low isotopic signatures (Supplementary Dataset? 4 ). The δ 13 C values of the hopanes are markedly lower than those observed in the co-occurring bulk organic matter (?24.0 to ?32.0‰; Fig.? 3a, c ) and n -alkanes ( n -C 21 : ?33.0 to ?38.1‰). Such 13 C-depleted hopanoids also appear in some modern/Holocene (e.g., Lake Rotsee, Switzerland 33 ) and Eocene (e.g., Green River Formation, USA 46 ) lake systems, where aerobic CH 4 oxidation by methanotrophic bacteria was prevalent in the water column. Here we conducted a survey of hopanoid δ 13 C values from 19 lakes (283 data points; δ 13 C hopanoid ranging from ?22.2 to ?71.9‰; Supplementary Fig.? 8 and Supplementary Dataset? 6 ). The data compilation (see Supplementary Note? 3 for data overview) suggests that hopanoid δ 13 C values below ?40‰ are indicative of a pronounced aerobic methanotroph contribution to these compounds (>10–20%; calculated from a C isotopic mass-balance approach; 33 see Methods).
Hopanoid δ 13 C values can be used to trace the temporal changes in aerobic CH 4 oxidation 33 , 42 , 44 , 45 . In this study, the consistently low hopane δ 13 C values (10‰ (from ca. ?46 to ?63‰ for C 30 hopane and from ca. ?43 to ?54‰ for C 29 hopane; Fig.? 3c ). The lowest values within the uppermost stratigraphic interval are among the most 13 C-depleted reported in the C 30 and C 29 hopanes for lacustrine systems (Supplementary Fig.? 8 ). These isotopic signatures indicate that substantially intensified CH 4 oxidation occurred in the water column, which closely coincided with elevated CH 4 production in the sediments, as indicated by a temporal increase in dolomite δ 13 C values (Fig.? 3b, c ). The combined evidence from both authigenic dolomite and molecular fossil (hopane) suggests that an intensification of the microbial CH 4 cycling occurred during the Artinskian age. Furthermore, active CH 4 cycling had a wide geographical distribution in paleo-Lake Junggar, with evidence of similar 13 C-depleted hopanes also documented in the Lucaogou shales from the adjacent Sangonghe section 22 and the Santanghu Basin 23 , hundreds of kilometers from the studied area (Fig.? 1 ). Based upon our estimated depositional duration (ca. 9?Myr) of the Lucaogou Formation, the intensified microbial CH 4 cycling persisted for at least ca. 3–5?Myr. To our knowledge, such a long-term dynamic of lacustrine CH 4 cycling in the Earth’s history has not been previously and directly revealed.
Positive feedback to Artinskian climate warming .
To investigate the relationship between temperature and microbial CH 4 cycling, we used the chemical index of alteration (CIA) 47 to reconstruct changes in the land surface temperature (LST; 18 , 48 see Methods). The collected samples were not affected by K-metasomatism, and their uniform Ti/Al ratios indicate no changes in provenance 48 (Supplementary Fig.? 9 and Supplementary Dataset? 7 ), and they thus provide a reliable record of climate variation. The CIA profiles show an increase from 50–55 in the lower member to 65–75 in the upper member, suggesting a rapid rise in the estimated LSTs from ~4?°C (Sakmarian) to ~14?°C (Artinskian; Fig.? 3d ). Overall, the pronounced progression toward higher CIA values, combined with the alternative chemical index of weathering (CIW; 49 Supplementary Fig.? 9 ), indicates a shift toward warmer conditions 14 , 18 , 48 . This record is consistent (within age uncertainties) with an independently derived CIA trend in a contemporaneous succession from the Karoo Basin of South Africa 19 , 21 (Fig.? 3e ). A cross-basin correlation revealed that a significant increase in CIA (temperature) globally began near the Sakmarian–Artinskian boundary (ca. 290?Ma) 18 . This major climate transition can be further corroborated by a coincident decrease in δ 18 O values from both low- and high-latitudinal fossil shells composed of low-Mg calcite 50 , 51 (Fig.? 3e ). Therefore, the elevated continental weathering in this study reflects a global climate warming signal (i.e., the Artinskian Warming Event 14 ), which developed contemporaneously with the intensification of CH 4 cycling in paleo-Lake Junggar (Fig.? 3 ).
Higher temperatures may have stimulated methanogenesis in lake sediments, supporting a temperature control on CH 4 cycling at the ecosystem level 8 , 42 , 52 . It has been proposed that the metabolic responses of methanogens are particularly sensitive to increases in temperature 8 , 52 . Since the predominant microbial methanogenesis occurred in the shallow sediment columns 53 , it would be expected that the increase in atmospheric temperature warmed the sediments and subsequently facilitated methanogenic activity. Additionally, under global warming, enhanced continental weathering (Fig.? 3d ) may have increased riverine nutrient influx and aquatic productivity in lakes, thereby resulting in increased substrate (e.g., acetate and H 2 /CO 2 ; ref. 6 ) availability for methanogenesis 8 . However, methanotrophy is known to have a more positive effect on substrate (i.e., CH 4 ) availability than temperature 8 , and the intensified CH 4 consumption observed in the top part of the Lucaogou Formation (mid-Artinskian) was almost certainly a response to an increased CH 4 substrate supply for methanotrophs (Fig.? 3 ).
The balance between methanogenesis and methanotrophy ultimately controlled the amount of CH 4 released into the atmosphere 6 , 8 , 9 . Nonetheless, if a warming-induced increase in CH 4 production exceeds the increase in CH 4 oxidation, an increase in net CH 4 emissions is expected, and this provides potential positive feedback to climate warming. Indeed, owing to the different temperature sensitivities of methanogens and methanotrophs 8 , 9 , warming would increase CH 4 emissions, which has been extensively observed in both modern freshwater ecosystems 7 , 52 , 54 and laboratory incubations 8 , 9 , 52 . For example, experimental warming of artificial ponds has suggested a disproportionate increase in methanogenesis over methanotrophy 9 . Although aerobic methanotrophs did oxidize more CH 4 , but not enough to offset the greater warming-induced CH 4 production 9 . Methane fluxes from lake ecosystems exhibit a temperature dependence 8 , 52 , 54 . The prevailing paradigm of the exponential response of CH 4 emissions to temperature 7 , 8 , 52 , 54 can be extrapolated to ancient lake systems, and the total CH 4 emissions from paleo-Lake Junggar could potentially have increased by several-fold in response to Artinskian climate warming. Applying the average CH 4 flux (total 31.6 Tg CH 4 yr ? 1 in areas spanning 1,330,264?km 2 ; i.e., 65?mg CH 4 m –2 d –1 ) 2 from modern lakes at similar latitudes to paleo-Lake Junggar (paleolatitude of 39–43°N) 13 , the flux was roughly estimated as 6.4 Tg CH 4 yr –1 (accounting for 5–28% of annual lake CH 4 emissions in the modern world 4 ), and a total amount of ~19,200 Gt CH 4 was emitted from this Early Permian lake (~270,000?km 2 ; ref. 11 ; herein conservatively calculated using 3?Myr).
Although there is only evidence for intensified CH 4 cycling in paleo-Lake Junggar (Fig.? 1 ), this still provides a useful analogue for similar environments having responses to Artinskian (Early Permian) climate warming. In this respect, several contemporaneous lake systems (see Fig.? 1 and Supplementary Dataset? 8 for the locations of these lakes and associated essential information) may also be CH 4 emission hotspots. However, accurate assessments of global CH 4 emissions require clear constraints relating to the contemporaneous lake area, distribution, and environmental factors, and these are beyond the scope of this study. Nonetheless, large-scale lacustrine CH 4 emissions would have acted as a positive feedback to Artinskian global warming and a critical mechanism for deriving carbon cycle perturbations. During this time period (after 290?Ma) 15 , the demise of the Late Paleozoic Ice Age (LPIA) was supported by a 6-fold drop in documented glacial deposits (Fig.? 3f ) 20 and the full deglaciation in south-central Gondwana by 282 Ma 19 , representing one of the most prominent and enigmatic climate transitions in the Earth’s Phanerozoic history. Previous studies have demonstrated that widespread deglaciation was synchronous with an increase in atmospheric p CO 2 (Fig.? 3f ) 15 derived from volcanic eruptions (e.g., Tarim, Panjal, and Zaduo large igneous provinces) 15 , 29 , and this provides an evidence regarding the strong link between CO 2 and glaciation. In addition to the contribution of CO 2 (refs. 15 , 16 , 17 ) and potential methane clathrate release 55 , our results suggest that the injection of the terrestrial greenhouse gas CH 4 into the atmosphere may have facilitated the demise of the LPIA and played a direct role in forcing the turnover from a long-lived icehouse to a greenhouse world.
In summary, this study investigated currently unexplored lacustrine ecosystem-level microbial CH 4 cycling records, including methanogenesis and methanotrophy, in pre-Cenozoic sedimentary archives. Our results suggest that sustained and intensified CH 4 cycling, as a response to Artinskian (Early Permian) climate warming, occurred in paleo-Lake Junggar. The release of the greenhouse gas CH 4 from large paleo-lakes to the atmosphere could have provided a direct positive feedback to ancient global warming, at least during the Early Permian, which should improve our understanding of its role in near-future climate change within a warming-but-glaciated world. .
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