Feedbacks and the coevolution of plants and atmospheric CO2

The coupled evolution of land plants, CO2, and climate over the last half billion years has maintained atmospheric CO2 concentrations within finite limits, indicating the involvement of a complex network of geophysiological feedbacks. But insight into this important regulatory network is extremely limited. Here we present a systems analysis of the physiological and geochemical processes involved, identifying new positive and negative feedbacks between plants and CO2 on geological time scales. Positive feedbacks accelerated falling CO2 concentrations during the evolution and diversification of terrestrial ecosystems in the Paleozoic and enhanced rising CO2 concentrations across the Triassic-Jurassic boundary during flood basalt eruptions. The existence of positive feedbacks reveals the unexpected destabilizing influence of the biota in climate regulation that led to environmental modifications accelerating rates of terrestrial plant and animal evolution in the Paleozoic.     

Mode change of millennial CO2 variability during the last glacial cycle associated with a bipolar marine carbon seesaw

Important elements of natural climate variations during the last ice age are abrupt temperature increases over Greenland and related warming and cooling periods over Antarctica. Records from Antarctic ice cores have shown that the global carbon cycle also plays a role in these changes. The available data shows that atmospheric CO₂ follows closely temperatures reconstructed from Antarctic ice cores during these variations. Here, we present new high-resolution CO₂ data from Antarctic ice cores, which cover the period between 115,000 and 38,000 y before present. Our measurements show that also smaller Antarctic warming events have an imprint in CO₂ concentrations. Moreover, they indicate that during Marine Isotope Stage (MIS) 5, the peak of millennial CO₂ variations lags the onset of Dansgaard/Oeschger warmings by 250 ± 190 y. During MIS 3, this lag increases significantly to 870 ± 90 y. Considerations of the ocean circulation suggest that the millennial variability associated with the Atlantic Meridional Overturning Circulation (AMOC) undergoes a mode change from MIS 5 to MIS 4 and 3. Ocean carbon inventory estimates imply that during MIS 3 additional carbon is derived from an extended mass of carbon-enriched Antarctic Bottom Water. The absence of such a carbon-enriched water mass in the North Atlantic during MIS 5 can explain the smaller amount of carbon released to the atmosphere after the Antarctic temperature maximum and, hence, the shorter lag. Our new data provides further constraints for transient coupled carbon cycle-climate simulations during the entire last glacial cycle.     

Disentangling the effects of CO2 and short-lived climate forcer mitigation

Anthropogenic global warming is driven by emissions of a wide variety of radiative forcers ranging from very short-lived climate forcers (SLCFs), like black carbon, to very long-lived, like CO₂. These species are often released from common sources and are therefore intricately linked. However, for reasons of simplification, this CO₂–SLCF linkage was often disregarded in long-term projections of earlier studies. Here we explicitly account for CO₂–SLCF linkages and show that the short- and long-term climate effects of many SLCF measures consistently become smaller in scenarios that keep warming to below 2 °C relative to preindustrial levels. Although long-term mitigation of methane and hydrofluorocarbons are integral parts of 2 °C scenarios, early action on these species mainly influences near-term temperatures and brings small benefits for limiting maximum warming relative to comparable reductions taking place later. Furthermore, we find that maximum 21st-century warming in 2 °C-consistent scenarios is largely unaffected by additional black-carbon-related measures because key emission sources are already phased-out through CO₂ mitigation. Our study demonstrates the importance of coherently considering CO₂–SLCF coevolutions. Failing to do so leads to strongly and consistently overestimating the effect of SLCF measures in climate stabilization scenarios. Our results reinforce that SLCF measures are to be considered complementary rather than a substitute for early and stringent CO₂ mitigation. Near-term SLCF measures do not allow for more time for CO₂ mitigation. We disentangle and resolve the distinct benefits across different species and therewith facilitate an integrated strategy for mitigating both short and long-term climate change.For about two decades, policy-makers have considered options to avoid dangerous anthropogenic interference with the climate system (1). So far, many countries support limiting warming to below a 2 °C temperature limit, but the required global mitigation action to achieve this has been limited (2–4). To inform policy-makers about options and challenges, the United Nations Environment Program (UNEP) published several reports over the past years on three interlinked aspects: climate stabilization and greenhouse gas (GHG) mitigation (3), short-lived climate forcers (SLCFs) and clean-air benefits (5, 6), and hydrofluorocarbons (7) (HFCs). We build here upon the insights of these reports (henceforth referred to as “Gap Report,” “SLCF Reports,” and “HFC Report,” respectively) to disentangle the joint effects of CO₂ and SLCF mitigation for limiting global warming. We evaluate the potential for limiting global-mean warming until 2100 and the rate of near-term warming, with a focus on 2 °C-consistent scenarios (Fig. 1). Reductions in CO₂ and SLCFs also provide important cobenefits like energy security (8), and local health and agricultural benefits (9–12), which fall outside the scope of this paper.Open in a separate windowFig. 1.Influence of SLCF-CO₂ linkages under varying CO₂ mitigation. (A) Global-mean surface temperature implications and interdependence of CO₂ (black), CH₄ (green), HFC (orange), BC-related (blue), and SO₂ mitigation (red). (B) The general effect of SLCF-CO₂ linkages. CO₂ paths show a world “with CO₂ mitigation” (32) and with “no CO₂ mitigation” (24). Early CH₄ mitigation is represented by the combined light and dark green area. HFC mitigation is shown for the lower end of the range assessed in this study. BC-related (and SO₂) measures show the difference between Case 6 and Case 2 (Case 4 and Case 2). Alternative cases are provided in SI Appendix, Fig. S1. Vertical dashed lines are time points relevant to Figs. 2 and ​and33.The main challenge in this exercise is the interdependence of coemitted climate forcers and the differences between their net forcing effects (13). For example, energy-related black carbon (BC) aerosols have an overall warming effect (14), whereas sulfate aerosols and some biomass-related BC emissions together with their coemitted species are cooling (13, 14). Because CO₂ and BC-related emissions often have common combustion sources (14), CO₂ mitigation will also influence the abundance of SLCFs. This linkage has already been well studied for other air pollutants (15, 16). Due to data limitations, the first studies that analyzed the mitigation potential of SLCFs (5, 6, 9, 17–19) did not account for these linkages in the long term and kept post-2030 SLCF forcing constant across a wide range of CO₂ paths. Alternatively, simple relationships between species were used (20). Such approaches, however, cannot guarantee that the long-term SLCF and CO₂ evolutions remain internally consistent. To provide an integrated view, we here account for this linkage and apply relationships (21) derived from detailed energy–environment–economy scenarios that explore various levels of air pollution control and track technological linkages between SCLF and CO₂ sources (8). Each CO₂ scenario in our analysis is thus associated with a consistent evolution of SLCFs at a specific level of pollution control stringency (see below). In policy discussions, methane (CH₄) and BC are often subsumed under the single term “short-lived climate pollutants” (SLCP) but in light of their different influence on the climate, as well as differing technological and policy instruments for mitigation, they are explicitly distinguished here.     

Contributions to accelerating atmospheric CO2 growth from economic activity, carbon intensity, and efficiency of natural sinks

The growth rate of atmospheric carbon dioxide (CO(2)), the largest human contributor to human-induced climate change, is increasing rapidly. Three processes contribute to this rapid increase. Two of these processes concern emissions. Recent growth of the world economy combined with an increase in its carbon intensity have led to rapid growth in fossil fuel CO(2) emissions since 2000: comparing the 1990s with 2000-2006, the emissions growth rate increased from 1.3% to 3.3% y(-1). The third process is indicated by increasing evidence (P = 0.89) for a long-term (50-year) increase in the airborne fraction (AF) of CO(2) emissions, implying a decline in the efficiency of CO(2) sinks on land and oceans in absorbing anthropogenic emissions. Since 2000, the contributions of these three factors to the increase in the atmospheric CO(2) growth rate have been approximately 65 +/- 16% from increasing global economic activity, 17 +/- 6% from the increasing carbon intensity of the global economy, and 18 +/- 15% from the increase in AF. An increasing AF is consistent with results of climate-carbon cycle models, but the magnitude of the observed signal appears larger than that estimated by models. All of these changes characterize a carbon cycle that is generating stronger-than-expected and sooner-than-expected climate forcing.     

Plant uptake of CO2 outpaces losses from permafrost and plant respiration on the Tibetan Plateau

High-latitude and high-altitude regions contain vast stores of permafrost carbon. Climate warming may result in the release of CO₂ from both the thawing of permafrost and accelerated autotrophic respiration, but it may also increase the fixation of CO₂ by plants, which could relieve or even offset the CO₂ losses. The Tibetan Plateau contains the largest area of alpine permafrost on Earth. However, the current status of the net CO₂ balance and feedbacks to warming remain unclear, given that the region has recently experienced an atmospheric warming rate of over 0.3 °C decade⁻¹. We examined 32 eddy covariance sites and found an unexpected net CO₂ sink during 2002 to 2020 (26 of the sites yielded a net CO₂ sink) that was four times the amount previously estimated. The CO₂ sink peaked at an altitude of roughly 4,000 m, with the sink at lower and higher altitudes limited by a low carbon use efficiency and a cold, dry climate, respectively. The fixation of CO₂ in summer is more dependent on temperature than the loss of CO₂ than it is in the winter months, especially at higher altitudes. Consistently, 16 manipulative experiments and 18 model simulations showed that the fixation of CO₂ by plants will outpace the loss of CO₂ under a wetting–warming climate until the 2090s (178 to 318 Tg C y⁻¹). We therefore suggest that there is a plant-dominated negative feedback to climate warming on the Tibetan Plateau.

High-latitude and high-altitude regions have a harsh, cold climate that favors the storage of carbon (C) in the entire soil column, including the permafrost layer (1). This permafrost contains the largest store of C in terrestrial ecosystems (roughly 1,320 Pg C) (2). These cold ecosystems have experienced a much greater rate of climate warming (0.3 °C decade⁻¹) than the 0.12 °C decade⁻¹ of warming across the global land surface as a whole (3). Warming can result in the release of CO₂ from the permafrost layer (4–6) but can also lead to increased fixation of CO₂ by plants (7–9), which can relieve or even offset the loss of CO₂ from permafrost.The frozen soils in the high-latitude Arctic have experienced a warming rate of the permafrost ground temperature of 0.39 °C decade⁻¹ (10). This has affected the permafrost reserves of C, as well as plant growth (7), and may also provide a feedback to climate warming, although the direction and magnitude of this feedback are still highly uncertain (5, 8, 9). Tundra ecosystems might be a net CO₂ source as a result of the differential amplification of the C cycle under warming conditions (11)—that is, winter CO₂ losses are more temperature-sensitive than the uptake of CO₂ by plants during the growing season. A recent study showed that the Pan-Arctic permafrost is a net CO₂ source of 630 Tg C y⁻¹ (12), albeit with strong uncertainties (range: 15 to 975 Tg C y⁻¹). By contrast, a synthesis study and model simulations reported that the Arctic can act as a CO₂ sink at the ecosystem level (8, 13, 14), given that the uptake of CO₂ by plants is accelerating as a result of Arctic greening. These contrasting results reveal the large uncertainties in feedbacks to the present-day warming climate across Arctic permafrost ecosystems.High-altitude mountains also have vast regions of permafrost (15), but less is known about the CO₂ balance in these regions and its variation in the Earth’s currently changing climate. The Tibetan Plateau ranges from <3,000 to 8,844 m in altitude and is the largest region of alpine permafrost on Earth, contributing 10% to the global store of permafrost C (SI Appendix, Fig. S1) (16, 17). Continuous permafrost covers 84.2 million hectares of the Tibetan Plateau, discontinuous permafrost covers 15.9 million hectares, and isolated permafrost covers 23.0 million hectares (16, 17). The Tibetan Plateau has experienced climate warming of 0.3 °C decade⁻¹ in the atmospheric temperature and an increase in precipitation since the 1960s (18). In situ observations and simulations both suggest a significant increase in the ground temperature of 0.4 °C decade⁻¹ over the last 50 y, causing extensive thawing of the permafrost (19–21). High-resolution observations from remote sensing satellites and unmanned aerial vehicles have shown an acceleration in thaw slumps and thermokarst lakes (22, 23), although on relatively small scales (a few hundred meters to several kilometers). The thawing of permafrost may have affected the regional CO₂ balance by causing large losses of CO₂. In situ observations at thaw slumps, incubation experiments, and numerical modeling have all shown that permafrost is highly vulnerable to warming on the Tibetan Plateau (24–26). A recent model simulation predicted an 8% loss of C from permafrost soils by the end of the present century (25). However, repeated sampling surveys in the 2000s and 2010s suggested there has been an accumulation of soil C in these permafrost regions, largely due to increased plant growth (6), which strongly contradicts the conclusions of model simulations.Many studies, especially those using remote sensing observations, have shown that the Tibetan Plateau is becoming greener under the currently warming and wetting climate and reduced density of grazing animals (27–30), suggesting a higher input of plant C to the soils. These contrasting views—that is, the model-based projection of the loss of CO₂ from permafrost and the inventory-based accumulation of soil C and satellite-based observations of increased vegetation—mean that it is unclear whether the CO₂ balance has been altered by the changing climate on the Tibetan Plateau. The size of the net CO₂ sink of the alpine ecosystems on the Tibetan Plateau is not yet fully understood, and the large-scale patterns of CO₂ flux are still unknown. It is therefore a research priority to investigate how the release of CO₂ from alpine permafrost might affect the CO₂ balance on the Tibetan Plateau.Eddy covariance observations, owing to their high temporal resolution and landscape-scale coverage, are an ideal approach to the direct measurement of the exchange of CO₂. Eddy covariance observations can therefore be used to assess the status and variation of the CO₂ balance of permafrost regions; however, historically, the establishment of eddy covariance sites on the Tibetan Plateau has been rare and synthesis efforts limited. In this paper, we present an eddy covariance dataset covering 32 sites across the main alpine ecosystems of the Tibetan Plateau (Fig. 1, SI Appendix, Figs. S2 and S3, and Datasets S1 and S2), ranging from 3,000 to over 5,000 m in altitude during the period 2003 to 2019. We then analyze the drivers of the spatial pattern, with specific emphasis on the altitude-dependent pattern given the importance of altitudinal variation to high mountains like the Tibetan Plateau. The eddy covariance dataset was used to constrain a process-based model and project future changes in the CO₂ balance under the Representative Concentration Pathways (RCP) 2.6 to 8.5 scenarios using temperature and precipitation projections from 18 Coupled Model Intercomparison Project Phase 5 (CMIP5) models. To provide ground-based evidence about the feedback of the CO₂ balance of the alpine permafrost regions to warming, we also explored how the exchange of CO₂ reacts to changes in climate by examining 16 manipulated experiments across the Tibetan Plateau. By connecting the eddy covariance datasets, the process-based model, and the experiments, we aim to understand whether the Tibetan Plateau currently acts, and will continue to act, as a CO₂ sink under a rapidly changing climate.Open in a separate windowFig. 1.Location of the 32 eddy covariance observation sites on the Tibetan Plateau. The inset map shows the location of the Tibetan Plateau (denoted by the black line). The characters in each solid circle are the name of each eddy covariance site: BT, Batang; DX, Dangxiong; DS, Dashalong; QM, Mt Qomolangma (Everest); GL, Guoluo; HB, Haibei; LQ, La. Qinghai; HY, Haiyan; FH, Fenghuoshan; LJ, Lijiang; MZ, Muztag; MD, Maduo; NM, NamCo; NQ, Naqu; SL, Shule; SZ, Shenzha; TG, Tanggula; YK, Yakou; ZG, Zoige.     

A CO2 greenhouse efficiently warmed the early Earth and decreased seawater 18O/16O before the onset of plate tectonics

The low ¹⁸O/¹⁶O stable isotope ratios (δ¹⁸O) of ancient chemical sediments imply ∼70 °C Archean oceans if the oxygen isotopic composition of seawater (sw) was similar to modern values. Models suggesting lower δ¹⁸O_sw of Archean seawater due to intense continental weathering and/or low degrees of hydrothermal alteration are inconsistent with the triple oxygen isotope composition (Δ’¹⁷O) of Precambrian cherts. We show that high CO₂ sequestration fluxes into the oceanic crust, associated with extensive silicification, lowered the δ¹⁸O_sw of seawater on the early Earth without affecting the Δ’¹⁷O. Hence, the controversial long-term trend of increasing δ¹⁸O in chemical sediments over Earth’s history partly reflects increasing δ¹⁸O_sw due to decreasing atmospheric pCO₂. We suggest that δ¹⁸O_sw increased from about −5‰ at 3.2 Ga to a new steady-state value close to −2‰ at 2.6 Ga, coinciding with a profound drop in pCO₂ that has been suggested for this time interval. Using the moderately low δ¹⁸O_sw values, a warm but not hot climate can be inferred from the δ¹⁸O of the most pristine chemical sediments. Our results are most consistent with a model in which the “faint young Sun” was efficiently counterbalanced by a high-pCO₂ greenhouse atmosphere before 3 Ga.

The amount of carbon that degassed from a solidifying magma ocean on the infant Earth 4.5 Ga ago was probably similar to the amount of CO₂ that is now present in the atmosphere of Venus (pCO₂ ∼90 bar) (1). Subsequently, dynamic carbon cycling between the early Earth’s atmosphere (atmospheric reservoir [R_A]), the ocean (R_O), and the oceanic crust (R_OC) reservoirs stabilized a primordial Earth–atmospheric pCO₂ to about 1.5 bar (2). Decreasing pCO₂ over the Earth’s history reflects the net transfer of large masses of carbon out of the atmosphere-ocean-oceanic crust (R_A+O+OC) system into the mantle and the emerging continental crust reservoir (R_CC), the latter providing a long-term sink for carbon in the form of inorganic carbon (carbonate rocks) and organic material (organic matter–rich shales, coal, oil, and gas) (1, 2).Early carbon-cycle models predicted high Archean pCO₂ to account for the faint young Sun paradox (2, 3), whereas direct pCO₂ estimates from the end of the Archean now imply much lower atmospheric pCO₂ [∼0.01 to 0.1 bar (4)]. However, a compilation of evidence from the sedimentary record implies much higher pCO₂ between 3.2 to 3.0 Ga compared to the 2.9 to 2.7 Ga interval (5). These authors state that even several bars pCO₂ are feasible between 3.2 to 3.0 Ga, which is not in conflict with much lower Neoarchean pCO₂ estimates (Fig. 1) or paleo-atmospheric pressure estimates at 2.7 Ga [<2 bar (6) and <0.5 bar (7)]. A fundamental drop in atmospheric CO₂ mixing ratio is also reflected in the observation that >3 Ga, Archean mafic crust (greenstones) is commonly characterized by very intense carbonatization and silicification that is unparalleled in their modern analogs (8–12). Such observations provide evidence for deep-time paleo-pCO₂ fluctuations with a drastic pCO₂ drop starting around 3 Ga ago (5, 13).Open in a separate windowFig. 1.Constraints on the R_A+O+OC size over time with implications for pCO₂. A illustrates the size of the R_A+O+OC carbon reservoir and its distribution between the oceanic crust (R_OC in gray), ocean water (R_O in dark blue), and the atmosphere (R_A in light blue). Transition of this carbon to the long-term R_CC and the mantle (green) decreases the size of the R_A+O+OC reservoir, which was still large at 3.2 Ga (see SI Appendix) and minimal at the onset of the global glaciations at 2.4 Ga. B shows two recent pH curves over Earth’s history to illustrate how the pH-dependent distribution of carbon between R_A and R_O may translate into pCO₂ at a given time. (C) The panel summarizes pCO₂ estimates [adopted from Catling and Zahnle (4)] and proposed pCO₂ evolution curves illustrated as dashed lines: a (2), b (18), and c (proposed here). Qualitative evidence to construct curve c comes from rare evidence for glaciers (44) and lower degrees of carbonatization of oceanic crust at 3.5 Ga compared to 3.2 Ga (13), suggesting a transient interval of somewhat low pCO₂ in the Paleoarchean. Low pCO₂ is indicated prior to the onset of cold climates later in the rock record. The black bar at 4.5 Ga is derived from carbon-flux arguments and the primordial carbon reservoir (1, 2). The black bar at 3.2 Ga applies the same carbon-flux arguments to translate the high carbonate content observed in the oceanic crust in Pilbara into tentative pCO₂ estimates (SI Appendix), which are most consistent with curve a (2).Today, carbonatization, which is the formation of carbonates during alteration of the oceanic crust, mainly occurs in relatively cool, off-axis hydrothermal systems over the first 20 Mya after crust formation at midocean ridges (14–16). Elevated degrees of carbonatization in oceanic crust from the Cretaceous and Jurassic are assigned to higher dissolved inorganic carbon at the time (14). Hence, higher pCO₂ (i.e., a larger R_A+O+OC) directly translates into higher degrees of carbonatization. While carbonate is mainly observed as vein fillings in the upper 300 m of oceanic crust today (14, 15), it extensively replaces glass and igneous minerals in Archean greenstones down to depths of 2 km below the ancient seafloor (8–10, 17). The amount of CO₂ fixed in 3.2-Ga-old oceanic crust from Pilbara, Australia is estimated at 1.2 × 10⁷ mol ⋅ m⁻² (±10%) (17)—a remarkable figure that is about two orders of magnitude more compared to today (SI Appendix). Much lower degrees of carbonatization in oceanic crust are already observed at 2.6 Ga, suggesting a drastic Mesoarchean drop in pCO₂ (13) (Fig. 1A).The qualitative evidence from the sediment record (5) and from the degree of carbonatization and silicification of oceanic crust (13) has not been included in proposed pCO₂ curves (4, 18) because the quantitative conversions into absolute pCO₂ require some assumptions (SI Appendix). Nevertheless, pCO₂ during the early Archean may have been as high as initially predicted by Kasting (2), followed by a pronounced Mesoarchean drop (13) to levels consistent with available paleo-pCO₂ estimates toward the Neoarchean (4) (stippled line “c” in Fig. 1C). Further decreasing pCO₂ toward modern levels partly reflects increasing ocean pH (Fig. 1) rather than a shrinking R_A+O+OC (SI Appendix). Hence, only small effects of carbonatization on the δ¹⁸O_sw value are expected for post-Archean seawater (16). Here, we focus on the very high carbonatization (8–10, 13) and silicification (11, 12) fluxes before the Mesoarchean pCO₂ drop, and we model the respective effects on ancient δ¹⁸O_sw.     

Rapid timescale for an oxic transition during the Great Oxidation Event and the instability of low atmospheric O2

The Great Oxidation Event (GOE), arguably the most important event to occur on Earth since the origin of life, marks the time when an oxygen-rich atmosphere first appeared. However, it is not known whether the change was abrupt and permanent or fitful and drawn out over tens or hundreds of millions of years. Here, we developed a one-dimensional time-dependent photochemical model to resolve time-dependent behavior of the chemically unstable transitional atmosphere as it responded to changes in biogenic forcing. When forced with step-wise changes in biogenic fluxes, transitions between anoxic and oxic atmospheres take between only 10² and 10⁵ y. Results also suggest that O₂ between ~10−8 and ~10−4 mixing ratio is unstable to plausible atmospheric perturbations. For example, when atmospheres with these O₂ concentrations experience fractional variations in the surface CH₄ flux comparable to those caused by modern Milankovich cycling, oxygen fluctuates between anoxic (~10−8) and oxic (~10−4) mixing ratios. Overall, our simulations are consistent with possible geologic evidence of unstable atmospheric O₂, after initial oxygenation, which could occasionally collapse from changes in biospheric or volcanic fluxes. Additionally, modeling favors mid-Proterozoic O₂ exceeding 10−4 to 10−3 mixing ratio; otherwise, O₂ would periodically fall below 10−7 mixing ratio, which would be inconsistent with post-GOE absence of sulfur isotope mass-independent fractionation.

Abundant atmospheric O₂ at 21% by volume is the most distinctive and consequential feature of Earth’s atmosphere. Produced by cyanobacteria, algae, and plants, O₂ is a clear sign of our biosphere that is detectable across interstellar space by telescopic spectroscopy (1). Oxygen permits aerobic respiration, the only known metabolism with sufficient energy yield to sustain complex animal life (2). However, for about the first half of Earth’s 4.5-billion-year-old history, the atmosphere had negligible O₂ (e.g., ref. 3). This changed ~2.4 billion years ago.The timing of the Great Oxidation Event (GOE) and the magnitude of atmospheric O₂ concentrations before and after the GOE can be constrained by the geologic record of sulfur isotopes in combination with photochemical models. Archean and earliest Proterozoic sedimentary minerals contain sulfur isotopes with characteristic mass-independent fractionation (MIF) which abruptly disappears 2.4 billion years ago (4). Sulfur MIF in marine sediments likely requires that atmospheric photochemistry produce elemental sulfur, S₈ (for explanation, see the introduction in ref. 5) (6, 7). Zahnle et al. (5) used a one-dimensional (1D) photochemical model to show that atmospheric S₈ production only occurs when atmospheric O₂ is below ~2×10−7 mixing ratio. An often cited threshold of 2×10−6 was from an earlier photochemical model that did not simulate atmospheres with surface O₂ mixing ratios between 2×10−6 and ~10−15 (6). Therefore, the disappearance of the sulfur isotope MIF signal at 2.4 Ga is strong evidence that O₂ first rose above 2×10−7 mixing ratio then.Geologic evidence may suggest that the GOE was not a single monotonic rise of oxygen but characterized by oscillations. Using U-Pb dating, Gumsley et al. (8) updated the chronology of sulfur isotope MIF in the stratigraphic record, finding evidence for two oxic-to-anoxic transitions between ~2.4 and ~2.3 Ga. More recently, Poulton et al. (9) report 2.3 Ga to 2.2 Ga marine sediments with sulfur isotopes consistent with approximately five oxic-to-anoxic transitions. Fluctuating O₂ levels coincide with three to four widespread glaciations, indicating extreme climate instability (10). Overall, geochemical evidence tentatively suggests that O₂ concentrations and climate were unstable for 200 million years until 2.2 Ga, which marks the most recent estimated timing of the permanent oxygenation of the atmosphere (9). However, interpretations of oscillating O₂ have been questioned (11). While the geologic evidence for the O₂ oscillations remains equivocal, the data have raised significant questions regarding the feasibility and timescales for Earth’s great oxidation. Some have argued that the oxygen-rich atmosphere is more stable than an oxygen-poor atmosphere (12), which favors a single rise of O₂ instead of O₂ oscillations.Evidence for O₂ instability and the time-dependent behavior of O₂ concentrations has not been reconciled with atmospheric photochemical models. All previous models treated the GOE as successive photochemical steady states (5, 6, 13–19). A photochemical steady state occurs when no atmospheric species changes concentration over time, because their production and loss from reactions and surface sources (e.g., volcanoes or biology) are balanced. Such steady-state calculations have been crucial for understanding the GOE by contextualizing sulfur isotope MIF observations (5, 6), and establishing the relationship between atmospheric O₂ concentrations and the degree to which O₃ blocks UV photons from Earth’s surface (i.e., O₃ shielding) (13, 15, 16), but they do not evaluate time-dependent changes and transient imbalances, or characteristic timescales.Several theories for the rise of O₂ suggest that it relied on a global redox titration over 10⁸ y to 10⁹ y involving oxidation of the upper mantle and/or crust, plausibly driven by hydrogen escape, which led to a tipping point where the source flux of O₂ exceeded a kinetically rapid O₂ sink from volcanic and metamorphic reductants (20–24). Beyond the tipping point, O₂ flooded the atmosphere, reaching a new, long-term balance limited by oxidative weathering.Here, we developed a time-dependent 1D photochemical model capable of investigating changes of O₂ at the tipping point itself over timescales of 10² y to 10⁵ y rather than the longer-term planetary changes which initiated the GOE. We simulate changing O₂ as a time-dependent evolution, in contrast to the steady-state approach used in previous studies (e.g., ref. 13), because O₂ can change on relatively rapid timescales that are not well characterized by steady states. With our model, we compute the time required for an anoxic-to-oxic atmospheric transition, and the time required for deoxygenation. Additionally, we investigate the stability of O₂ concentrations against perturbations to surface gas fluxes produced by biology. Finally, we use our model results to better constrain O₂ levels and stability during the GOE (starting at ~2.4 Ga), and during the mid-Proterozoic eon (1.8 Ga to 0.8 Ga).     

Changes in plasma levels of adrenaline, noradrenaline, glucose, lactate and CO2 in the green turtle, Chelonia mydas, during peak period of nesting

Plasma concentrations of stress hormones [adrenaline (ADR), noradrenaline (NR)], lactate, glucose and CO2 were monitored during peak nesting period (May-October) at different phases of nesting in the green turtle, Chelonia mydas. These include, emergence from sea, excavating body and nest chambers, oviposition, covering and camouflaging the nest and then returning to sea. Turtles that completed all phases of nesting including oviposition before returning to sea were considered "successful" turtles, while those that completed all phases but failed to lay their eggs were "unsuccessful". Blood samples were taken from the cervical sinus within 5min of capture to avoid stress due to handling. The turtles were usually sampled for blood between 20:00 and 1:00h of nesting time to ensure uniformity in the sampling. Plasma ADR and NR values were highly significant (P<0.001) in successful turtles over emergence, excavating and unsuccessful turtles. Plasma glucose levels remained stable throughout the nesting phases while lactate levels were significantly higher in successful turtles over the other phases (P<0.05) which signifies anaerobic metabolism during nesting. Plasma CO2 values were negatively correlated with ADR and NR (r=-0.258, P=0.03; r=-0.304, P=0.010), respectively. Hematocrit was significantly higher in successful phase (P<0.05) compared to other phases, and this may signify a higher degree of stress in successful turtles. Body temperature were significantly lower (P<0.005) in the excavating phase compared to the other three phases. Overall, body temperatures were lower than sand temperatures around the nest, which may indicate a behavioral thermoregulation used by the turtles during nesting. This information will be of value to the ongoing conservation program at Ras Al-Hadd Reserve in the Sultanate of Oman.     

Effect of Samples Size on the Water Removal and Shrinkage of Eucalyptus urophylla × E. grandis Wood during Supercritical CO2 Dewatering

Eucalyptus urophydis E. grandis green wood with different lengths were dewatered using CO₂ that was cyclically alternated between the supercritical fluid and gas phases. The results indicate that shorter specimens can be dewatered to below the fiber saturation point (FSP). There was no significant difference in the dewatering rate between the specimens of 20 and 50 mm in length. The dewatering was faster when the moisture content (MC) was over the FSP, leading to a greater gradient and a non-uniform distribution of moisture. The MC distributions in all specimens had no clear differences between in tangential and radial directions. Supercritical CO₂ dewatering generated a different moisture gradient than conventional kiln drying. Most water was dewatered from the end-grain section of the wood along the fiber direction, but a small amount of water was also removed in the transverse directions. There was no deformation in the specimens when the MC was above the FSP.