Introduction — Carbon dioxide in Earth’s atmosphere
Continuous atmospheric CO2 measurements initiated at Mauna Loa Observatory in 1958, commonly presented as the Keeling Curve, record mole‑fraction in micromoles per mole (ppm) and show an uninterrupted rise from 1958 through 2023. On a molar basis CO2 reached about 427 ppm (0.0427%) by 2024, equivalent to roughly 3,341 gigatonnes of CO2 in the atmosphere; this concentration is approximately 50% higher than the ~280 ppm that prevailed throughout the 10,000 years before the mid‑18th century, and the modern rise is attributed to human activities.
Although CO2 is a trace gas, it is central to the greenhouse effect, the terrestrial and oceanic carbon cycles, and photosynthesis, and is one of the three principal long‑lived greenhouse gases. The dominant contemporary source of the increase is fossil‑fuel combustion, with substantial additional contributions from cement manufacture, land‑use change (including deforestation), and biomass burning. Higher atmospheric CO2, together with other long‑lived greenhouse gases such as methane, increases the atmosphere’s capacity to absorb and re‑emit infrared radiation, thereby raising global mean surface temperature (global warming) and driving ocean acidification. Elevated CO2 also directly affects plant physiology through the CO2 fertilization effect, enhancing photosynthetic uptake in some species, while triggering a suite of ecological and societal impacts mediated by climate change.
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The radiative influence of CO2 follows from its molecular spectroscopy: the CO2 molecule has two principal infrared‑active vibrational modes relevant to Earth’s radiation budget—an asymmetric stretch near 4.26 μm (≈2,347 cm−1) and a bending mode near 14.99 μm (≈667 cm−1). Earth’s thermal emission is strongest in the infrared band roughly between 200 and 2,500 cm−1, so these CO2 absorption bands overlap the outgoing terrestrial spectrum. Absorption of outgoing longwave radiation at these wavelengths traps energy in the lower atmosphere and surface layers while reducing upward energy flux, which also tends to cool the upper atmosphere.
Current atmospheric CO2 is the highest in about 14 million years, yet geological records demonstrate a much wider natural range over deep time. Reconstructions across the Phanerozoic show marked variability—examples include concentrations near 4,000 ppm in the Cambrian (~500 Ma), approximate 2,000 ppm peaks in the Devonian (~400 Ma), pronounced Triassic highs (~220–200 Ma), and lows near 180 ppm during the Quaternary glaciations of the last two million years—illustrating that present values, while large in the recent geologic context, fall within a broader long‑term envelope of Earth system variability.
Global accounting by the Global Carbon Project for 1850–2019 indicates that about two‑thirds of cumulative anthropogenic “excess” CO2 resulted from fossil‑fuel combustion and associated industrial processes. Of those fossil‑fuel emissions, slightly less than half accumulated in the atmosphere, so that the persistent atmospheric contribution attributable to fossil fuels over this period is therefore somewhat under one‑third of the total excess anthropogenic CO2. The remainder of fossil‑fuel emissions was taken up by natural sinks—predominantly the oceans and the terrestrial biosphere—producing geographically significant outcomes such as increased oceanic carbon uptake with attendant acidification and greater carbon storage in some land regions. These figures are global, cumulative aggregates that reflect the long history of industrialization and spatially uneven fossil‑fuel use; the fraction of emissions retained in the atmosphere is the primary driver of the observed long‑term rise in atmospheric CO2 and its climatic impacts.
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Current situation
Since the onset of the Industrial Revolution atmospheric carbon dioxide concentrations have increased substantially, a principal driver of observed global warming and a major cause of ongoing ocean acidification as the oceans absorb additional CO2. Continuous monitoring by the U.S. National Oceanic and Atmospheric Administration reports a seasonally adjusted global mean of 422.17 ppm for October 2023. By contrast, the Holocene pre‑industrial background over the roughly 10,000 years before the mid‑18th century was about 280 ppm, providing the baseline for assessing modern change.
For mass‑equivalence, each 1 ppm of atmospheric CO2 corresponds to roughly 2.13 gigatonnes of carbon (≈7.82 Gt CO2). Since 1850 human activities have emitted on the order of 2,650 Gt CO2 in total, with current annual emissions near 42 Gt CO2 yr−1; of those cumulative emissions about 1,050 Gt CO2 currently remain in the atmosphere after uptake by oceans and terrestrial sinks. A significant fraction of emitted fossil carbon is effectively long‑lived: model projections indicate that approximately 20–35% of the carbon already released will persist as elevated atmospheric CO2 for many thousands of years after anthropogenic emissions decline.
Paleoclimate and instrumental records reviewed through 2021 show that the present rates of increase of the principal greenhouse gases—CO2, CH4 and N2O—are unprecedented over at least the last 800,000 years, indicating a rate of change that far exceeds natural variability on glacial–interglacial timescales.
Annual and regional variability in atmospheric CO2 is dominated by the seasonal rhythm of terrestrial vegetation. In the Northern Hemisphere spring and summer, photosynthetic uptake by extensive mid‑latitude plant biomass produces a substantial drawdown of atmospheric CO2; between May and September this intra‑annual net uptake amounts to roughly 6–7 ppm (on the order of 50 Gt). As vegetation growth terminates and biomass decays through autumn and winter, CO2 concentrations rebound by approximately 8–9 ppm. The seasonal extrema are phase‑locked to the NH growing cycle: concentrations commonly peak in May and decline to a seasonal minimum around October.
The Northern Hemisphere imposes the principal imprint on the global seasonal signal because it contains markedly more land area and vegetative mass concentrated in the 30°–60° latitude band; this latitudinal distribution amplifies NH seasonal carbon exchange relative to the Southern Hemisphere. Spatial variability of CO2 is most pronounced near the surface, where local sources and sinks—ecosystems, soils, combustion—generate regional differences. Vertical gradients decline with altitude, so concentrations aloft are substantially more homogeneous.
Human environments accentuate these patterns at local scales: urban areas systematically exhibit higher CO2 than surrounding regional backgrounds due to concentrated emissions, and poorly ventilated indoor spaces can reach CO2 levels an order of magnitude above ambient outdoor values. These surface‑proximate heterogeneities must be accounted for when interpreting regional measurements and attributing fluxes.
Instrumental measurements over the past decade document a clear and accelerating increase in atmospheric carbon dioxide. Global-mean CO2 was rising at roughly 2 ppm yr−1 by 2009, with the rate described as accelerating, implying larger annual additions to the atmospheric reservoir and a steadily increasing global baseline. Regional records capture this rise unevenly: the Arctic registered daily concentrations near 400 ppm in June 2012, and the long-term Mauna Loa record recorded its first daily average above 400 ppm on 10 May 2013, demonstrating both spatial variability and the complementary roles of polar monitoring and the Mauna Loa baseline in tracking change. Crossing the ~400 ppm threshold was significant historically and paleoenvironmentally—exceeding the full instrumental era and reaching levels probably not experienced since the Pliocene (on the order of millions of years). Measurements continued upward thereafter, reaching about 410 ppm by 2018, consistent with the ongoing trend of increasing atmospheric CO2.
Measurement techniques
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Atmospheric carbon dioxide is measured continuously and through discrete sampling at a global network of sites; records from 2008–2017 confirm pronounced, recurring seasonal oscillations superimposed on a rising background concentration, together with a systematic interhemispheric offset. Continuous instrumental monitoring—most notably the Mauna Loa record begun in 1958—provides one of the longest, internally consistent time series and serves as a reference for seasonal cycles and hemispheric differences. The methodological foundation for reproducible long‑term monitoring was laid by flask‑sample measurements developed by C. D. Keeling in the 1950s, which established sampling protocols and reference standards still used for calibration and cross‑station comparison.
CO2 concentrations are reported in parts per million by volume (ppmv, often written ppm(v) or simply ppm). To convert to parts per million by mass (ppmm or ppm(m)), multiply ppmv by the ratio of the molar mass of CO2 to the molar mass of dry air (44.01/28.96 ≈ 1.52). Using consistent units and conversion factors is essential for rigorous geographical and temporal comparison across datasets.
Spatial and temporal patterns in measured CO2 reflect the interplay of surface exchange and atmospheric transport. The larger seasonal amplitude in the northern hemisphere is attributed primarily to its greater terrestrial land area and seasonal vegetation dynamics—strong biospheric uptake during the northern growing season and release during dormancy—whereas the southern hemisphere’s signal is damped by its larger ocean fraction and weaker seasonal biospheric exchange. Large‑scale circulation and interhemispheric mixing attenuate but do not eliminate these contrasts, producing persistent but evolving differences in phase and amplitude between hemispheres.
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Coordinated measurement networks and publicly available datasets enable spatial attribution of regional sources and sinks, cross‑validation of long‑term trends (for example against the Mauna Loa benchmark), and consistent comparison across stations. Robust interpretation of CO2 observations therefore requires standardized units and conversion, reproducible sampling and calibration procedures, and analysis framed by land–ocean distribution, biospheric seasonality, and atmospheric transport processes.
Data networks
A coordinated constellation of surface measurement networks underpins global observations of atmospheric carbon dioxide, employing both discrete flask sampling and continuous in situ instruments. Prominent contributors include NOAA/ESRL, the World Data Centre for Greenhouse Gases (WDCGG) and national databases such as France’s RAMCES; together these networks supply the primary observational records used to characterize atmospheric CO2 concentrations.
In the United States, the NOAA/ESRL Baseline Observatory Network and the Scripps Institution of Oceanography maintain long-term measurement series that are archived and distributed via the Carbon Dioxide Information Analysis Center (CDIAC) at Oak Ridge National Laboratory, providing organized repositories for baseline, coastal and remote site data. Internationally, the WDCGG operates within the Global Atmosphere Watch (GAW) system and—hosted by the Japan Meteorological Agency—aggregates contributions from many countries, while RAMCES (affiliated with the Institut Pierre‑Simon Laplace) represents a national node feeding the broader observational system.
The two principal measurement techniques—flask sampling, which yields discrete, laboratory-calibrated samples at lower temporal frequency, and continuous in situ measurements, which produce high-temporal-resolution concentration time series—are complementary but produce variability in sampling frequency and temporal coverage across sites. Because individual networks differ in instrument types, temporal span and spatial distribution, raw observations exhibit gaps and uneven coverage that complicate direct synthesis.
To address these discontinuities, integrated data products merge observations from multiple networks into coherent, gap-filled datasets suitable for temporal trend analysis, model forcing and inter-network comparison. GLOBALVIEW-CO2 exemplifies this approach, combining heterogeneous measurements into a consistent global CO2 product that mitigates spatial and temporal sparseness and facilitates robust scientific and modelling applications.
Analytical methods to investigate sources of CO2
The isotopic composition of carbon in atmospheric CO2 provides a diagnostic means to separate fossil‑fuel emissions from contemporary biological fluxes: carbon released from long‑buried fuels carries a different set of carbon‑isotope ratios than carbon exchanged with the living biosphere, and these differences can be used quantitatively to estimate the anthropogenic fraction of observed CO2 changes.
Spatial and compositional signals in the atmosphere reinforce this attribution. Measured CO2 concentrations are persistently higher in the Northern Hemisphere because most people and fossil‑fuel combustion occur there, and the north‑south concentration difference has grown with rising emissions. At the same time, oxidation of fossil carbon consumes molecular oxygen, producing a detectable decline in atmospheric O2 that is directly linked to combustion of fossil fuels.
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Taken together, isotope measurements, interhemispheric CO2 gradients, and concurrent O2 trends form a coherent, geographically informed toolkit for attributing changes in atmospheric composition to human activities. These complementary signals—distinct isotopic fingerprints of fossil carbon, an increasing Northern‑dominant CO2 gradient, and falling O2—reflect both the spatial concentration and temporal rise of emissions and constitute the empirical foundation for global carbon‑cycle monitoring and attribution studies.
Anthropogenic CO2 emissions
Since the industrial revolution, human activities have become the dominant driver of the modern rise in atmospheric carbon dioxide. The extraction and combustion of geologic fossil carbon—coal, oil, gas and associated industrial processes such as cement production—have added large quantities of CO2 that natural sinks cannot fully absorb, producing a net accumulation in the atmosphere (by May 2022 atmospheric CO2 was about 50% above pre‑industrial levels). In recent years fossil‑fuel combustion released on the order of 9 GtC per year (~30+ GtCO2 yr−1; 2019), a scale of input responsible for the observed upward trend in atmospheric concentration.
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Fossil‑fuel combustion (including cement manufacture) remains the primary anthropogenic source of CO2, with land‑use change and deforestation the second major source. Emissions from fossil fuels plus cement rose markedly from 6.15 GtC in 1990 to 9.14 GtC in 2010 (the latter equivalent to ~33.5 GtCO2, or roughly a 4.3 ppm addition to the atmosphere), while estimated emissions from land‑use change declined over the same interval (from about 1.45 GtC in 1990 to 0.87 GtC in 2010), altering the relative contributions of different sectors.
Natural sinks—principally terrestrial vegetation and the oceans—remove roughly half of the CO2 emitted by fossil‑fuel combustion, so approximately 50% of those emissions persist in the atmosphere and contribute to long‑term radiative forcing. Over longer historical intervals the pace of emissions accelerated dramatically: fossil‑fuel burning released roughly 12 GtC between 1751 and 1900 but about 380 GtC between 1901 and 2013, reflecting exponential growth during the industrial and post‑industrial eras.
Cumulative emissions have been spatially uneven, with the United States, China and Russia among the largest historical emitters since 1850 and thus major determinants of the global distribution of anthropogenic CO2. Emissions are also highly unequal at the individual level: in 2021 the International Energy Agency estimated that the top 1% of emitters each had footprints exceeding 50 tCO2 (more than 1,000 times those of the bottom 1%), whereas the global average energy‑related carbon footprint was roughly 4.7 tCO2 per person.
Greenhouse effect
The greenhouse effect arises because atmospheric gases are largely transparent to incoming shortwave solar radiation but absorb and re‑emit outgoing longwave (infrared) radiation from Earth’s surface; this retention of thermal energy raises surface temperatures relative to a no‑atmosphere state. Spectrally, different gases absorb at distinct infrared wavelengths according to their longwave absorption coefficients, producing characteristic features in the outgoing spectrum.
Carbon dioxide exhibits a pronounced absorption band centered near 667 cm−1 (≈15 μm) and in that spectral region is a much stronger absorber of outgoing longwave radiation than water vapor. Because this band lies in a portion of the infrared where additional absorption can still occur, increases in CO2 measurably reduce the flux of thermal radiation escaping to space and thus increase net radiative forcing.
Water vapor contributes the largest single share of the present greenhouse effect—commonly estimated between about 36 and 70 percent—but its concentration depends strongly on temperature, making it principally an amplifying feedback rather than a directly controlled forcing. By contrast, CO2 is the most climatically and geopolitically significant gas subject to human control because anthropogenic emissions have raised its atmospheric abundance since the pre‑industrial era (circa 1750). For example, in 2013 the CO2 increase was estimated to account for roughly 1.82 W m−2 of the total 2.63 W m−2 change in radiative forcing, or about 70 percent of that increment.
The theoretical link between higher atmospheric CO2 and warmer surface temperatures dates back to Svante Arrhenius (1896). Contemporary understanding explains the continued effectiveness of added CO2 by the existence of non‑saturated absorption windows—infrared wavelengths where current CO2 does not completely block outgoing radiation—so that further CO2 increases raise atmospheric opacity and radiative forcing. Over time, this enhanced forcing alters Earth’s energy balance and drives changes in surface temperature, atmospheric circulation, and the interacting feedbacks of the climate system.
Carbon cycle
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The carbon-cycle schematic quantifies exchanges among the atmosphere, land and oceans in billions of metric tons of carbon per year, using color coding to distinguish natural fluxes (yellow), anthropogenic contributions (red) and stored reservoirs (white). Atmospheric CO2 functions as the principal coupling medium between the biosphere, lithosphere and hydrosphere: it is removed by photosynthesis and by chemical precipitation of carbonates (e.g., limestone formation) and returned by biological respiration and by acid-driven dissolution of carbonate deposits.
Two interacting cycles operate on markedly different timescales. The fast carbon cycle comprises rapid exchanges between living and recently dead organic matter (vegetation, soils, surface-ocean biology) and the atmosphere. The slow cycle involves long-term transfers among the atmosphere, surface and deep ocean, soils, sedimentary rocks and volcanism; atmospheric CO2 links the fast and slow reservoirs and mediates transfer between them.
Natural sources of atmospheric CO2 include volcanic outgassing, wildfire combustion of organic matter and aerobic respiration. Contemporary volcanic emissions are small in comparison with other fluxes, on the order of 130–230 megatonnes CO2 per year. Human-caused emissions derive primarily from fossil-fuel combustion and industrial processes such as cement manufacture; annual anthropogenic flows have been systematically reported since 1960 in units of gigatonnes of carbon per year (Gt C yr−1).
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Terrestrial fluxes illustrate the large magnitude and near balance of natural exchanges: decay and fire-related releases from forests, grasslands and other vegetation amount to roughly 436 Gt CO2 yr−1 (≈119 Gt C yr−1), while uptake by regrowth and primary production removes about 451 Gt CO2 yr−1 (≈123 Gt C yr−1), so that terrestrial uptake slightly exceeds terrestrial release. Historically, the terrestrial biosphere shifted roles: from the pre‑industrial era through about 1940 it acted as a net source of CO2—largely because of land‑use change—whereas after ~1940 it became a net sink as vegetation regrowth and other processes removed more CO2 than land change released.
Carbon cycles continuously among atmospheric CO2, live and dead vegetation, soils, the ocean surface layer and the deep ocean; the differing sizes and exchange timescales of these reservoirs produce contrasting short‑term and long‑term carbon storage. A commonly used, simplified representation of atmospheric persistence is the Bern-derived impulse response function (Joos et al.), which gives the fraction f(t) of a one‑time CO2 pulse remaining in the atmosphere after t years as
f(t) = 0.217 + 0.259 exp(−t/172.9) + 0.338 exp(−t/18.51) + 0.186 exp(−t/1.186).
This formulation implies that about 21.7% of an emitted CO2 pulse remains on very long timescales in the absence of engineered or geological removal; deliberate sequestration could, however, alter that long‑term fraction.
Data and diagrams employ two related unit conventions—billions of metric tons of carbon per year in the schematic fluxes and equivalent gigatonnes of carbon per year when reporting historical anthropogenic flows—so that natural flux magnitudes, human inputs and reservoir stocks can be compared consistently across time and space.
Oceanic carbon cycle
The global ocean constitutes the largest reservoir of Earth’s inorganic carbon, predominantly as dissolved bicarbonate (HCO3−) and carbonate (CO32−) ions. This dissolved inorganic carbon (DIC) pool, far exceeding the atmospheric CO2 inventory, largely derives from chemical weathering of rock and dissolution of carbonates that transform CO2 into stable aqueous species, thereby controlling marine carbon speciation.
Between 1850 and 2022 the ocean absorbed roughly 26% of cumulative anthropogenic CO2 emissions. This uptake has moderated atmospheric accumulation but has altered seawater chemistry—driving ocean acidification—and varies spatially according to regional uptake efficiency and circulation patterns.
Progressive changes in the marine carbonate system shift DIC speciation: increased dissolution of carbonate minerals tends to raise bicarbonate concentrations while decreasing or leaving unchanged carbonate ion levels, which in turn raises the proportion of un‑ionized carbonic acid (H2CO3) and dissolved molecular CO2. Because air–sea CO2 exchange responds to the concentration of dissolved CO2 in surface waters and to temperature-dependent gas solubility (Henry’s law), higher dissolved CO2 together with warmer sea surfaces elevate the equilibrium partial pressure of CO2 over the ocean. That state reduces future oceanic uptake capacity and can promote outgassing relative to colder or less carbon‑loaded conditions.
Southern Ocean circulation exerts a key dynamical control on these fluxes. Observations and models (e.g., recent analysis in Science Advances, 2025) indicate that an accelerated Antarctic Circumpolar Current enhances upwelling of deep, isotopically light waters around Antarctica; surfacing of such waters releases CO2 to the atmosphere. The interplay of altered carbonate chemistry, temperature‑dependent solubility, and circulation‑driven upwelling—notably ACC‑induced upwelling—therefore constitutes a plausible positive feedback on warming, increasing uncertainty in future ocean CO2 uptake and generating important regional heterogeneity with global climate implications.
The figure synthesizes physical drivers responsible for observed global warming, distinguishing realized forcings from projected future contributions and explicitly excluding the additional warming potential of continued emissions of long‑lived agents such as CO2; thus it represents past and present forcing contributions rather than forward projections. Each bar quantifies an individual driver’s contribution to observed warming and is accompanied by whiskers that denote uncertainty bounds, indicating that attribution magnitudes carry statistical uncertainty and thereby constrain confidence in the estimated global contributions. Rising atmospheric CO2 produces a direct radiative forcing that elevates global mean temperatures; although the mechanism operates at the planetary scale, its warming effects are spatially heterogeneous across regions and climates, and the graphic isolates this realized warming from any further warming that would result from future CO2 accumulation. Increased atmospheric CO2 also drives ocean acidification through dissolution into seawater and consequent pH decline, a global process with regionally variable impacts on marine chemistry and ecosystems that disproportionately affects calcifying organisms and associated food webs. Finally, higher CO2 concentrations can enhance photosynthetic rates and potential productivity in some terrestrial ecosystems and crops (the CO2 fertilization effect), but the magnitude and geographic distribution of this benefit are limited by local constraints such as water and nutrient availability and other environmental factors.
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Temperature rise on land
Global surface temperature (GST) combines land and ocean surface observations and functions as a principal metric of planetary warming, providing one of several independent lines of evidence for changes in Earth’s climate system. Observed centennial-scale temperature changes are consistent with the expected response to increasing atmospheric greenhouse gases: perturbations of the radiative balance drive a climate-system response manifested as warming.
Multiple independent datasets estimate that the global average combined land and ocean surface temperature rose by 1.09 °C (dataset range: 0.95–1.20 °C) between 1850–1900 and 2011–2020. This figure is a global mean that integrates conditions across terrestrial and marine environments and therefore represents a planet-wide shift rather than a local or regional anomaly. The rate of warming since the 1970s exceeds that of any other 50‑year interval in at least the last 2,000 years, indicating an acceleration of warming in the late 20th and early 21st centuries. The agreement among independent datasets, the quantified global increase, and the recent acceleration together substantiate the scientific consensus that the observed large-scale warming is driven by human-induced increases in greenhouse-gas concentrations.
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The global ocean reached its warmest state observed by instrumental records in 2022, with ocean heat content (OHC) surpassing the previous peak recorded in 2021. Because it integrates heat throughout the water column, OHC is the most reliable indicator of long‑term ocean warming and shows an accelerating accumulation of heat in recent decades.
This persistent warming reflects an underlying planetary energy imbalance: more incoming than outgoing energy at the top of the atmosphere. Rising atmospheric concentrations of greenhouse gases are the principal driver of that imbalance, making continued ocean warming a locked‑in consequence of the enhanced greenhouse effect.
Surface measurements provide a complementary benchmark: sea‑surface temperatures increased by roughly 0.68–1.01 °C between the pre‑industrial era and the 2011–2020 decade. Heat uptake is not uniform; the Southern Ocean accounts for a disproportionate share of global OHC gain and thus exerts a dominant influence on how heat is distributed worldwide. Regional records illustrate this concentration—between the 1950s and 1980s the Antarctic sector of the Southern Ocean warmed by about 0.17 °C, a rate nearly double the global ocean’s mid‑20th‑century increase.
Ocean acidification refers to the progressive, long‑term reduction in seawater pH caused principally by uptake of anthropogenic CO2. Global mean sea‑surface pH has fallen from about 8.15 to 8.05 between 1950 and 2020, concurrent with rising atmospheric CO2 concentrations that have exceeded ~422 ppm by 2024 and driving increased CO2 flux into the ocean surface.
Chemically, dissolved CO2 reacts with seawater to form carbonic acid (H2CO3), which dissociates to bicarbonate (HCO3−) and releases hydrogen ions (H+); the increased H+ concentration is the direct cause of lower pH. Because pH is a logarithmic measure of H+, relatively small numerical changes imply large chemical shifts: a 0.1‑unit drop in mean ocean pH corresponds to roughly a 26% rise in hydrogen‑ion concentration. Despite this decline, average seawater remains alkaline (pH > 8).
Ecologically, reductions in carbonate ion availability and carbonate saturation state associated with acidification undermine the ability of calcifying organisms—such as corals, mollusks and some plankton—to precipitate calcium carbonate, impairing growth, structural integrity and ecosystem services. The magnitude and biological consequences of acidification are not uniform: surface pH and carbonate saturation vary with depth, latitude and temperature, with colder, high‑latitude waters typically absorbing more CO2 and exhibiting stronger pH and saturation declines.
Regional and local patterns are further shaped by physical and biogeochemical processes. Ocean currents, upwelling (which brings CO2‑rich, low‑pH deep water to the surface), freshwater inputs from large rivers (which alter alkalinity and buffering capacity), and sea‑ice coverage (which influences gas exchange and stratification) all modulate the atmosphere–ocean CO2 exchange and local acidification. Anthropogenic and natural inputs of nitrogen and sulfur from fossil‑fuel combustion and agriculture additionally interact with nutrient supply, biological productivity and carbonate chemistry, producing complex spatial variability in acidification impacts.
CO2 fertilization refers to the direct physiological effects of elevated atmospheric carbon dioxide on plants: higher CO2 concentrations tend to increase photosynthetic carbon assimilation while reducing stomatal conductance and hence transpiration. The magnitude of these responses varies widely among species and is mediated by local conditions—air and soil temperatures, water availability and nutrient supply all modulate how strongly individual plants and communities respond.
At the ecosystem scale, enhanced photosynthetic rates under elevated CO2 often raise net primary productivity (NPP) and have been implicated in long‑term “greening” of vegetated lands since the 1980s. Remote‑sensing and flux‑based analyses attribute a substantial fraction of observed gains in gross primary productivity (GPP)—on the order of tens of percent, ~44% in some assessments since the 2000s—to CO2‑driven physiological change. However, higher instantaneous carbon uptake does not necessarily translate into proportional increases in whole‑plant growth or durable carbon storage: allocation patterns, nutrient and water limitations, and increased respiration or turnover can decouple photosynthetic gains from net biomass accumulation and long‑term sequestration.
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Understanding and projecting the CO2 fertilization effect relies on large‑scale simulation frameworks (Earth System Models, Land System Models and Dynamic Global Vegetation Models). These tools are essential for attributing productivity trends and testing scenarios, but key ecosystem processes that determine the magnitude and persistence of fertilization—such as nutrient feedbacks, changes in tissue turnover, and vegetation demography—remain uncertain and are difficult to represent robustly.
Consequently, although terrestrial vegetation has historically taken up carbon and thereby moderated increases in atmospheric CO2, the capacity of the land sink to offset anthropogenic emissions is limited. Under plausible future emission trajectories, plant responses to CO2 are unlikely to substantially lower atmospheric CO2 concentrations over the coming century. Simple expectations that the tropics should dominate CO2‑driven uptake, given their high baseline productivity, have not been unambiguously borne out by observations; realized sequestration depends critically on how tropical forests respond to climate change and on land‑use outcomes such as conservation versus deforestation.
Anthropogenic increases in atmospheric CO2 have produced a measurable downward displacement of stratospheric pressure and geopotential surfaces, estimated at roughly 400 m since about 1980. This signal denotes a change in the vertical placement (and effectively the thickness insofar as pressure surfaces move) of the stratospheric layer—not its disappearance—and should be understood in the context of the stratosphere’s nominal extent (~10–50 km, varying with latitude and season).
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The contraction is explained by radiative physics: additional CO2 enhances infrared emission to space from the stratosphere, producing net radiative cooling of that layer even as CO2 warms the troposphere. Cooler stratospheric temperatures reduce the atmospheric scale height (the exponential decay length of pressure with altitude), causing pressure/geopotential surfaces to reside at lower geometric altitudes. The observed 400 m change is a multi‑decadal, globally detectable mean that nonetheless exhibits latitude-, season- and circulation-dependent departures from the global average.
This vertical restructuring of atmospheric mass has practical consequences for technological and geodetic systems. Orbital prediction and mission planning for low Earth orbit (LEO) spacecraft rely on atmospheric density profiles; systematic lowering of stratospheric geopotential heights alters those profiles and thus affects drag estimates, orbital decay rates, stationkeeping fuel budgets, collision-avoidance calculations and lifetime projections. Precise positioning and timing applications (including geodetic reference frames) depend on atmospheric delay models: although the troposphere and ionosphere dominate signal delay, changes in stratospheric temperature and pressure modify refractivity and integrated delay terms sufficiently that updated atmospheric corrections are necessary to avoid small but systematic biases.
Electromagnetic propagation is also sensitive to the altered vertical structure. Refractive-index gradients and stability regimes that govern tropospheric ducting, scattering and long‑range propagation of HF, VHF and microwave signals will be affected by the shifted stratification, requiring adjustments in propagation modelling to preserve reliability of communications and remote sensing retrievals.
Maintaining operational accuracy therefore demands sustained monitoring (radiosondes, limb- and nadir-sounding satellites, and reanalysis products), routine revision of atmospheric-correction algorithms, and explicit incorporation of evolving stratospheric structure into satellite-tracking, navigation, radio‑propagation and remote-sensing toolchains.
Indirect effects and impacts
The contemporary climate system shows a sustained warming trend accompanied by shifts in precipitation regimes and an increase in extreme weather frequency and intensity. These changes reflect perturbations to the coupled atmosphere–ocean–land energy balance and produce spatially heterogeneous consequences for regional climates, hydrology and storm behaviour. Anthropogenic emissions—chiefly from combustion of fossil fuels and land‑use change such as deforestation—have raised atmospheric concentrations of CO2, methane and other greenhouse gases, amplifying radiative forcing and driving the observed climatic departures from preindustrial conditions.
On land, warming drives a suite of interrelated physical and ecological responses: higher incidence and severity of wildfires, progressive thawing of permafrost in high‑latitude and high‑altitude regions, and expansion or intensification of arid and semi‑arid conditions. These processes reorganize ecosystem structure and function, perturb biogeochemical cycles (including mobilization of previously stored carbon), and undermine the suitability of landscapes for some human uses. Many impacts are nonlinear; crossing critical thresholds or “tipping points”—for example abrupt permafrost carbon release or large‑scale biome transitions—can produce effectively irreversible changes, amplifying long‑term ecosystem loss and imposing substantial costs for adaptation. Societal responses to these risks have emerged at global and local scales through mitigation, adaptation planning, policy advocacy and direct action aimed at reducing emissions and protecting vulnerable environments.
Oceans play a central moderating role but are themselves undergoing profound change. The ocean has absorbed the majority of excess heat added to the climate system, which has warmed the upper ocean, increased the frequency and intensity of marine heatwaves, and redistributed thermal energy vertically and horizontally. Warming drives sea‑level rise both through thermal expansion of seawater and by adding meltwater from glaciers and ice sheets; concomitant declines in sea ice extent alter coastal and polar environments and feedback onto regional climate. Continued uptake of anthropogenic CO2 is lowering seawater pH (ocean acidification) and, together with warming‑related processes, is reducing dissolved oxygen in many regions (deoxygenation), with adverse consequences for marine organisms and biogeochemical functioning. Enhanced surface warming and freshwater inputs have increased ocean stratification, reducing vertical mixing of heat, oxygen and nutrients, altering productivity and habitat suitability, and potentially influencing large‑scale circulation systems (for example, a weakening of the Atlantic meridional overturning circulation). The ocean currently sequesters roughly one quarter of human‑emitted CO2, providing a partial buffer against atmospheric accumulation but at the cost of progressive chemical and biological changes in marine environments.
Approaches for reducing CO2 concentrations
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High‑resolution numerical modelling of atmospheric carbon over the period 1 September 2014 to 31 August 2015, using vertical exaggeration (~40×) to expose three‑dimensional flow patterns, demonstrates the spatial and temporal complexity of carbon transport in the atmosphere. Such models are instrumental for understanding how emitted CO2 is redistributed and for evaluating the likely efficacy and deployment constraints of mitigation measures.
Carbon dioxide differs from shorter‑lived greenhouse gases in its climatic persistence: a substantial fraction of emitted CO2 remains active in the climate system for centuries to millennia, rendering its climatic effects effectively near‑irreversible on thousand‑year timescales after emissions cease. On decadal to centennial timescales, surface air temperature after cessation of emissions is controlled by an energy‑balance interplay between reduced radiative forcing (from declining atmospheric CO2) and a simultaneous reduction in ocean heat uptake as ocean and atmosphere temperatures equilibrate. Because ocean warming continues even after emissions stop, sea surface and subsurface temperatures will rise for some time, driving continued thermal expansion and additional sea level rise independent of ongoing ice melt.
Given these system inertias, simply halting emissions would not produce a rapid return to pre‑industrial temperatures; achieving faster temperature declines requires active interventions to remove heat or decrease greenhouse forcing. Practically, this implies deliberate removal of atmospheric CO2 (carbon dioxide removal, CDR) or geoengineering measures that alter Earth’s energy balance.
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CDR denotes human actions that extract CO2 from the atmosphere and sequester it durably in geological formations, terrestrial carbon pools, ocean reservoirs, or long‑lived products. Increasingly integrated into climate policy, CDR complements deep emissions reductions: reaching “net zero” relies on first minimizing emissions and then applying CDR to offset residual, hard‑to‑abate sources (e.g., certain agricultural and industrial emissions). A range of CDR methods has been proposed to fulfil this role, and their selection and scale will depend on technical feasibility, permanence of storage, and environmental and socio‑economic considerations.
Concentrations in the geologic past
Reconstructions of atmospheric CO2 extend from deep time (hundreds of millions of years) down to the most recent 40,000 years, and they demonstrate both large long‑term swings and rapid short‑term changes. On multimillion‑year scales, recent syntheses (2023) indicate that contemporary CO2 levels may be the highest in roughly 14 million years; the IPCC Sixth Assessment Report similarly notes CO2 magnitudes during the mid‑Pliocene warm period (about 3–3.3 Ma) that are comparable to today and often serve as an analogue for the climate responses expected under sustained elevated greenhouse gas concentrations.
Carbon dioxide has been a dominant control on Earth’s surface temperature throughout its 4.54 billion‑year history. High early CO2 (and likely enhanced methane) are central to resolving the “faint young Sun” paradox—geological evidence for persistent liquid water despite a substantially dimmer young Sun. Earth’s atmosphere evolved in stages: an initial primary atmosphere soon after accretion, followed by a secondary atmosphere produced largely by volcanic outgassing and late‑bombardment volatiles, in which nitrogen and CO2 were abundant. Much of the CO2 emitted in these early stages was rapidly taken up by the oceans and precipitated as carbonate sediments. Estimates for CO2 partial pressure during very early Earth reach values as high as ~1,000 kPa (10 bar) because biological sinks such as photosynthetic carbon fixation were not yet established. The emergence of oxygenic photosynthesis in cyanobacteria precipitated the Great Oxidation Event (~2.4 Ga), introducing free oxygen and driving the transition toward the modern, oxygen‑rich atmosphere.
Throughout the Phanerozoic and Quaternary, atmospheric CO2 has varied widely: values on the order of ~4,000 ppm in the Cambrian (≈500 Ma) declined over geological time to minima near ~180 ppm during the last major glacial intervals (around 20 ka). Quaternary glacial–interglacial cycles repeatedly oscillated between about 180 ppm in deep glacials and ~280 ppm in interglacials, underscoring the tight coupling between greenhouse gas concentrations, global temperature, and ice volume. Superimposed on this natural variability is the recent, unprecedented rise in CO2: reconstructions from the Last Glacial Maximum to the present show that the modern rate of atmospheric CO2 increase far exceeds any rate observed during the most recent deglaciation, implying a novel trajectory for climate relative to past natural changes.
On geological timescales the atmospheric partial pressure of CO2 is controlled by a balance among slow, large‑scale geochemical fluxes at Earth’s surface and within the crust and mantle: burial of organic carbon in sediments, chemical weathering of silicate rocks, and volcanic degassing of subsurface carbon. Burial and silicate weathering act as long‑term sinks, whereas volcanic and tectonic release of buried carbon supply the atmosphere; persistent, modest imbalances among these fluxes, integrated over tens to hundreds of millions of years, have driven a gradual net decline in atmospheric CO2 through deep time.
Extending this perspective to billion‑year timescales, the downward trend is expected to continue because the most energetic, episodic injections of previously stored carbon—large igneous provinces, plume‑related volcanism and similar tectono‑volcanic events—should become progressively rarer. That reduction in extreme volcanic carbon release is linked to secular cooling of the mantle and the progressive depletion of radiogenic heat sources that sustain magmatism and long‑term tectonic activity.
Crucially, the characteristic rates of organic burial, silicate weathering and volcanic degassing are vanishingly slow relative to human and even glacial‑interglacial timescales. Averaged over tens of millions to billions of years the geologic carbon cycle is effectively sink‑dominated, but its slow kinetics mean these long‑term sinks cannot meaningfully moderate or counteract atmospheric CO2 changes occurring on centennial to millennial timescales.
Photosynthesis in the geologic past
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Atmospheric CO2 has been a dominant selective agent throughout Earth’s history, shaping which metabolic strategies and organismal groups prevailed in terrestrial and aquatic ecosystems. Early phototrophs are thought to have relied on inorganic reductants such as H2 or H2S rather than water, conducting anoxygenic photosynthesis that did not release O2 and thereby influenced primordial biogeochemical cycles. The later emergence of cyanobacteria introduced oxygenic photosynthesis and an attendant rise in free oxygen—the so‑called oxygen catastrophe—which fundamentally altered Earth’s redox state and opened ecological space for the subsequent diversification of aerobic, metabolically complex life.
In more recent geologic intervals, persistently lowered atmospheric CO2 (notably concentrations below ~600 ppm) appears to have imposed directional selection favoring the C4 photosynthetic pathway over the ancestral C3 pathway. By biochemically concentrating CO2 at the site of Rubisco and reducing photorespiration, the C4 mechanism conferred a competitive advantage under declining CO2 and associated environmental stresses, precipitating a major expansion of C4‑dominated vegetation between roughly 7 and 5 million years ago. Under present atmospheric pressure, many vascular plants exhibit critical lower limits for effective carbon fixation—on the order of ~150–200 ppm—below which conventional photosynthesis fails. Nevertheless, certain microbes can fix carbon at substantially lower ambient CO2, permitting persistence and continued microbially mediated carbon cycling under extreme low‑CO2 conditions.
Measuring ancient‑Earth CO2 concentration
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The most direct record of pre‑instrumental atmospheric CO2 arises from air occlusions in polar ice sheets. Gas bubbles preserved within Antarctic and Greenland cores supply physically trapped samples of ancient atmospheres for chemical analysis; classic Antarctic records (e.g., Vostok) show a close covariation of CO2 and reconstructed temperature through successive glacial–interglacial cycles. East Antarctic cores provide the longest continuous chronology, extending to about 800,000 years and yielding direct CO2 measurements for that interval. Over the last 800 kyr, ice‑core data indicate glacial minima near 180–210 ppm and interglacial maxima typically around 280–300 ppm; broadly, pre‑industrial values lay between ~180 and 280 ppm, with the Holocene remaining relatively stable at ~260–280 ppm until the Industrial Era. Atmospheric CO2 has since risen by roughly 35% from ~280 ppm in the 19th–early 20th century to ~387 ppm by 2009, a departure from the natural variability documented in the ice‑core record.
Ice‑core gas measurements must be interpreted with methodological constraints in mind. Air is incorporated gradually as firn pores close and undergoes layer mixing during burial, so each analytical value represents a multi‑decadal to multicentennial average rather than an instantaneous annual concentration. Ice cores also preserve concurrent changes in other greenhouse gases (notably CH4) that covary with temperature and CO2 across glacial cycles. However, the direct gas record terminates at the age limit of the ice archive, necessitating independent approaches for earlier intervals.
For times beyond the ice‑core window, multiple proxies and model reconstructions are employed: boron and carbon isotope ratios in marine carbonates, stomatal density and index on fossil leaves, and molecular biomarkers such as phytane derived from chlorophyll breakdown. Phytane and other biomarkers can provide continuous estimates that bridge gaps between archives; combined proxy–model evidence implies far larger long‑term CO2 fluctuations in deep time, with concentrations around 500 million years ago plausibly an order of magnitude higher than today. Some stomatal‑based reconstructions suggesting >300 ppm between ~10,000 and 7,000 years ago remain contested and may reflect calibration, regional bias, or contamination rather than global atmospheric signals.
720–400 Ma
Geochemical models and proxy records yield divergent reconstructions of atmospheric CO2 for the interval spanning the late Neoproterozoic through the early–middle Palaeozoic. Model-based estimates before the mid-Ordovician (≈450 Ma) often indicate very high CO2—on the order of thousands of parts per million—whereas organic biomarker (phytane) proxies for the Ordovician suggest substantially lower concentrations, roughly 300–700 ppm. More generally, CO2 throughout much of the Phanerozoic is inferred to have exceeded present levels: Mesozoic values are commonly estimated at ~4–6 times modern, and early Palaeozoic values at ~10–15 times modern. These reconstructions carry considerable uncertainty and reflect methodological differences between geochemical modelling and proxy approaches.
Key climatic transitions within this interval are linked to both tectonic and biological processes. In the Neoproterozoic an extended, intermittent near‑global glaciation (the Snowball Earth events) persisted for tens of millions of years and ended abruptly at ≈635 Ma when sustained volcanic outgassing is thought to have elevated atmospheric CO2 to extremely high concentrations (on the order of 10^5 ppm, ~12% by volume). The rapid greenhouse forcing that followed produced swift deglaciation and prolific carbonate precipitation, with inferred post‑glacial limestone accumulation potentially reaching tens of centimetres per year. Later, through the early to middle Palaeozoic, elevated CO2 levels declined markedly by the middle Devonian (~400 Ma) in association with the colonization and expansion of terrestrial vegetation; the proliferation of land plants both drew down CO2 and thereafter established dynamic source–sink feedbacks that helped stabilise the global carbon cycle and climate. The end of the Cryogenian and the attendant greenhouse rebound are also implicated in setting environmental conditions that preceded the Ediacaran biotic assemblages and may have contributed to the subsequent radiation of metazoans.
After roughly 60 million years ago atmospheric carbon dioxide entered a persistent Cenozoic decline. Geochemical proxy records document this long-term fall in greenhouse-gas concentrations and provide empirical constraints on key intervals of climatic change.
Around the Eocene–Oligocene transition (≈34 Ma), proxy-based reconstructions place atmospheric CO2 on the order of 10^2–10^3 ppm (estimates near 760 ppm), a period that coincides with major biotic turnovers and the initial development of the Antarctic continental ice sheet. Subsequent reduction in CO2—crossing a reconstructed climatic tipping range near ~600 ppm—has been identified as the principal forcing that promoted Antarctic glaciation and the establishment of a persistent ice sheet.
CO2 continued to decline through the Miocene; by ~20 Ma concentrations inferred from multiple geochemical proxies had fallen below ~300 ppm, marking markedly lower greenhouse forcing than in earlier Paleogene times. Prolonged low atmospheric CO2 also affected terrestrial ecosystems: reduced pCO2 created selective pressures that favored the physiological advantages of C4 photosynthesis, and the relative abundance and geographic spread of C4-dominated taxa rose substantially between about 7 and 5 Ma.