Sustainability & Environment (Agri-Food Systems) · Climate Science & Agriculture · also Physical Science of Climate Change, Climate Change Science, Global Climate Change
The physical basis of climate change as assessed by the IPCC: the greenhouse effect, radiative forcing, the global carbon cycle, feedbacks, climate sensitivity, and the evidence for observed change. It goes well beyond the 'climate change and livestock' unit of veterinary environmental hygiene.
In the exams: GATE ES Section 8 (greenhouse gases, global warming); UGC NET Environmental Sciences Unit 10; GARP SCR 'Foundations of Climate Change'; UPSC CSE environment
Climate system science is the physical basis of climate change: how the Sun's energy is absorbed, moved around and returned to space by the atmosphere, ocean, ice, land and living things, and how human activity has altered that flow. It covers the greenhouse effect and the Earth's energy budget, radiative forcing and the metrics used to compare greenhouse gases, the carbon, methane and nitrous oxide cycles, aerosols, feedbacks and climate sensitivity, the observed record of change, how scientists attribute that change to its causes, the link between cumulative carbon dioxide emissions and warming, tipping points, and the changing ocean. It is taught from the Intergovernmental Panel on Climate Change (IPCC) assessments, above all the Working Group I contribution to the Sixth Assessment Report (AR6), and this manual uses AR6 numbers unless it says otherwise.
The subject matters in practice because almost every later decision in sustainability rests on it. A carbon accountant converting methane to carbon dioxide equivalent is choosing a metric from this subject; a net-zero target is a statement about carbon budgets; a crop insurer pricing drought risk needs to know what attribution science can and cannot say; a regulator setting a methane rule needs to know why methane cuts slow warming within decades while carbon dioxide cuts only stop it from rising. Getting the physics wrong produces confident but wrong policy.
Worldwide, the subject is examined at three levels. Entrance and postgraduate papers (GATE Environmental Science and Engineering Section 8, CUET-PG, ICAR PhD Environmental Sciences, UGC NET Environmental Sciences Unit 10, CSIR NET atmospheric science) ask for the greenhouse effect, global warming potentials, carbon cycle pools, energy-balance arithmetic and IPCC facts. Civil-service papers such as UPSC CSE test the same facts in applied, current-affairs form. Professional credentials, notably the GARP Sustainability and Climate Risk certificate ('Foundations of Climate Change') and the CFA Institute Sustainable Investing Certificate, test whether a practitioner can read IPCC findings correctly: carbon budgets, confidence language, sensitivity ranges and tipping-point risks. This manual is written for all three.
Climate is the statistical description of weather over a long period, conventionally 30 years, in terms of the mean and variability of temperature, precipitation, wind and other quantities. The climate system is the coupled whole that produces it, and it has five components. The atmosphere is the fast, thin layer: it holds little heat and adjusts to a change in its boundary conditions in days to weeks. The hydrosphere, mainly the ocean, is the great heat store: the heat capacity of the whole atmosphere is matched by only the top few metres of sea water. The upper mixed layer, typically tens of metres deep, responds over years; the deep ocean over centuries to millennia. The cryosphere covers sea ice, snow, glaciers, permafrost and the Greenland and Antarctic ice sheets; sea ice and snow respond in seasons to years, while ice sheets take millennia to reach a new balance. The land surface sets albedo, roughness and evaporation. The biosphere, on land and in the ocean, exchanges carbon, water and energy with the atmosphere and links the physical system to the carbon cycle. These components interact. The ocean takes up heat and carbon dioxide and releases them on its own schedule; ice changes albedo; vegetation changes evaporation and carbon storage. The spread of response times is the key to much of the subject: it explains why the surface responds to a forcing within years but keeps warming slowly for centuries, why sea level keeps rising long after temperature stabilises, and why commitments made now are felt by people not yet born.
Climate changes for two kinds of reason. External forcings are imposed from outside the climate system as modelled: changes in solar output, volcanic eruptions that load the stratosphere with sulphate, changes in the Earth's orbit, and human changes to greenhouse gases, aerosols and land cover. Internal variability is generated inside the coupled system with no change in forcing at all, because the atmosphere and ocean exchange heat and momentum chaotically. The best known mode of internal variability is the El Niño–Southern Oscillation (ENSO), a coupled ocean–atmosphere oscillation of the tropical Pacific with a period of roughly two to seven years. El Niño years raise global mean surface temperature for a year or so; La Niña years lower it. Other modes include the North Atlantic Oscillation, the Pacific Decadal Variability and the Indian Ocean Dipole, which shapes monsoon rainfall in South Asia and East Africa. A single year or a decade can therefore be warmer or cooler than the forced trend, which is why trends are judged over decades. Over geological time the dominant pacemaker of ice ages has been orbital forcing, described by Milankovitch theory: changes in eccentricity (main periods near 100,000 and 400,000 years), obliquity or axial tilt (about 41,000 years, between roughly 22.1° and 24.5°) and precession (about 19,000 to 23,000 years). These change the seasonal and latitudinal distribution of sunlight far more than its annual global total, and the glacial–interglacial swing is amplified by feedbacks from ice albedo and carbon dioxide, which varied between about 180 parts per million (ppm) in glacials and about 280 ppm in interglacials over the last 800,000 years, as recorded in Antarctic ice cores. Palaeoclimate is therefore evidence for the strength of feedbacks, not a reason to doubt present-day attribution.
The Sun delivers about 1361 W m⁻² at the top of the atmosphere on a surface facing it at the Earth's mean distance; this is the total solar irradiance, often still called the solar constant. Because the Earth intercepts sunlight on a disc of area πR² but has a surface area of 4πR², the average incoming flux over the whole globe is one quarter of this, about 340 W m⁻². Roughly 29–30 per cent is reflected straight back by clouds, aerosols, bright surfaces and the atmosphere; this fraction is the planetary albedo. About 240 W m⁻² is absorbed. A body in balance must radiate what it absorbs. The Stefan–Boltzmann law gives the flux radiated by a black body as σT⁴, with σ = 5.67 × 10⁻⁸ W m⁻² K⁻⁴. Setting σTₑ⁴ = S(1 − α)/4 gives the effective emission temperature Tₑ. With S = 1361 W m⁻² and α = 0.30, Tₑ is about 255 K, or −18 °C. The observed global mean surface temperature is about 288 K (15 °C). The difference of about 33 K is the natural greenhouse effect. It is not a human creation: without it the Earth would be largely frozen. The greenhouse effect works because the atmosphere is nearly transparent to incoming shortwave sunlight but absorbs and re-emits outgoing longwave (infrared) radiation. Water vapour is the largest contributor to the natural effect, followed by carbon dioxide, with clouds, ozone, methane and nitrous oxide contributing smaller parts. Nitrogen and oxygen, the bulk of the air, have no significant infrared absorption because their symmetric two-atom molecules have no changing dipole moment when they vibrate. A greenhouse gas absorbs in particular bands: carbon dioxide's strong band near 15 μm sits close to the peak of the Earth's emission spectrum, while the atmospheric window between about 8 and 13 μm lets much surface radiation escape directly.
The clearest way to understand the enhanced greenhouse effect is through emission height. Most of the infrared radiation that escapes to space comes not from the surface but from an effective emission level in the mid-troposphere, where the air is at about 255 K. Because temperature falls with height in the troposphere, at a lapse rate of roughly 6.5 K per kilometre on average, the surface is warmer than the emission level. Adding a greenhouse gas raises the emission level to colder air, which radiates less; the planet then absorbs more than it emits until the whole troposphere and surface warm enough to restore balance. This picture also explains why greenhouse warming is accompanied by cooling of the stratosphere, where added carbon dioxide radiates more energy to space. Global energy-flow estimates in AR6 put the surface's emitted longwave radiation near 398 W m⁻², balanced in large part by about 342 W m⁻² of downward longwave 'back radiation' from the atmosphere. The surface also loses energy by evaporation (latent heat, roughly 80 W m⁻²) and by conduction and convection (sensible heat, roughly 20 W m⁻²). The outgoing longwave radiation at the top of the atmosphere is about 239 W m⁻². These figures carry uncertainties of several W m⁻² each. The difference between absorbed sunlight and outgoing longwave radiation is the Earth's energy imbalance. AR6 assessed that it grew from about 0.50 W m⁻² over 1971–2006 to about 0.79 W m⁻² over 2006–2018. About 91 per cent of the extra energy has gone into the ocean, about 5 per cent into warming land, about 3 per cent into melting ice, and about 1 per cent into warming the atmosphere. The imbalance is the most direct measure of how far the climate is from equilibrium.
Radiative forcing quantifies how strongly a driver pushes the energy budget. It is the change in net downward radiative flux, in W m⁻², caused by a change in a driver such as a gas concentration, with the climate response held back in a defined way. The definition has evolved. Instantaneous radiative forcing allows nothing to adjust. Stratospherically adjusted radiative forcing lets the stratosphere reach a new radiative equilibrium, which happens within months, while the surface and troposphere are held fixed. Effective radiative forcing (ERF), introduced as the preferred measure in AR5 and used throughout AR6, also lets fast adjustments in the troposphere, clouds and land surface happen while global mean surface temperature is held fixed. ERF predicts the eventual temperature response more consistently across very different drivers, such as carbon dioxide, absorbing aerosols and solar change. For carbon dioxide, a widely taught simplified expression is ΔF = 5.35 ln(C/C₀) W m⁻², from Myhre and colleagues (1998); the logarithm reflects saturation of the band centre, so each doubling adds roughly the same forcing, about 3.7 W m⁻² by this formula. AR6 assessed the ERF of a doubling of carbon dioxide as 3.93 W m⁻², higher because of updated spectroscopy and the inclusion of adjustments. Exam answers should state which value they use. AR6 assessed the total anthropogenic ERF in 2019 relative to 1750 at 2.72 W m⁻² (very likely range 1.96 to 3.48). Carbon dioxide contributed about 2.16 W m⁻², methane about 0.54, nitrous oxide about 0.21 and halocarbons about 0.41; tropospheric ozone added a positive term; aerosols contributed about −1.1 W m⁻² with a wide range. Aerosols are the largest single source of uncertainty in total forcing.
Inventories and targets need one number for many gases, so emission metrics convert a mass of another gas into a mass of carbon dioxide equivalent (CO2-e). The global warming potential (GWP) is the time-integrated radiative forcing from a pulse emission of 1 kg of a gas over a horizon H, divided by the same integral for 1 kg of carbon dioxide. GWP100, with H = 100 years, is the metric adopted for reporting under the Paris Agreement's transparency framework. AR6 values of GWP100 are 27.0 for methane from non-fossil sources and 29.8 for fossil methane (fossil methane also adds carbon dioxide when it oxidises), 273 for nitrous oxide, and values in the tens of thousands for sulphur hexafluoride and some fluorinated gases. Earlier assessments gave different numbers: AR4 gave 25 for methane and 298 for nitrous oxide, used for the Kyoto Protocol's second period; AR5 gave 28 and 265 without climate–carbon feedbacks. A carbon footprint is only reproducible if the GWP set is named. The choice of horizon matters greatly for a short-lived gas. Methane's perturbation lifetime is about 12 years, so its GWP20 (above 80 in AR6) is roughly three times its GWP100. The global temperature-change potential (GTP) is an end-point metric: the temperature change at year H from a pulse of the gas, relative to carbon dioxide. Because a methane pulse has largely decayed after 100 years, methane's GTP100 is far smaller than its GWP100, in single digits. GWP and GTP treat a tonne of methane as a fixed number of tonnes of carbon dioxide, yet the two gases act differently: warming from carbon dioxide depends on cumulative emissions, while warming from a short-lived gas depends mainly on its current rate of emission. GWP* (GWP-star) was developed by Allen, Cain and colleagues to capture this. It converts a change in the emission rate of a short-lived gas into a 'warming-equivalent' emission. In the form of Cain and colleagues (2019), E* = GWP_H × (r × ΔE/Δt × H + s × E), with Δt = 20 years, H = 100, r = 0.75 and s = 0.25; this simplifies to 4.0E(t) − 3.75E(t − 20) times GWP100. A later recalibration by Smith and colleagues (2021) gives 4.53E(t) − 4.25E(t − 20). GWP* shows that a steady methane source adds little new warming, while a rising one adds a great deal. It is a modelling and analysis tool, not the inventory standard, and its use to claim 'climate neutrality' for a historical emitter is contested because it treats the existing emission level as a baseline.
Carbon moves between reservoirs on very different time scales. Pools are given in petagrams of carbon (PgC; 1 PgC = 1 GtC = 10¹⁵ g). The atmosphere held about 590 PgC before industrialisation and holds about 870–880 PgC today; one part per million of carbon dioxide corresponds to about 2.12 PgC, or about 7.8 Gt of carbon dioxide (multiply carbon by 44/12 to get carbon dioxide). The ocean is by far the largest active pool, about 38,000 PgC, mostly as dissolved inorganic carbon in the intermediate and deep ocean; the surface ocean exchanges with the air within a year or so. Land vegetation holds roughly 450–650 PgC and soils roughly 1500–2400 PgC, depending on the depth counted. Permafrost soils of the northern high latitudes hold roughly 1400–1600 PgC, much of it frozen. Fossil fuel reserves are of the order of 1000 PgC. Ranges are wide because inventories differ in depth and method. The fast carbon cycle involves photosynthesis and respiration on land, air–sea gas exchange, and the ocean's biological and solubility pumps. Gross primary production on land is about 120 PgC per year, almost all returned by respiration and fire. The slow carbon cycle, operating over tens of thousands to millions of years, involves rock weathering, carbonate sedimentation and volcanic outgassing; silicate weathering is the thermostat that regulates carbon dioxide over geological time but is far too slow to remove human emissions on human time scales. The ocean takes up carbon dioxide by dissolution, which shifts the carbonate equilibria: CO2 + H2O ⇌ H2CO3 ⇌ H⁺ + HCO3⁻ ⇌ 2H⁺ + CO3²⁻. Most dissolved inorganic carbon is bicarbonate. Uptake consumes carbonate ions, and the Revelle (buffer) factor, typically about 10 at the surface and higher in cold water, measures how a given fractional rise in seawater carbon dioxide produces only a much smaller fractional rise in total dissolved carbon. As carbonate is used up, the Revelle factor rises and the ocean's buffer capacity falls.
The Global Carbon Project publishes an annual budget that partitions human emissions. For 2010–2019 it estimated fossil carbon dioxide emissions (including cement) of about 9.6 PgC per year and emissions from land-use change of about 1.6 PgC per year. Of these, about 5.1 PgC per year accumulated in the atmosphere, the ocean sink took about 2.5 PgC per year and the land sink about 3.4 PgC per year, with a small budget imbalance. Sinks are not constant: the land sink varies strongly from year to year, weakening in El Niño years. The airborne fraction is the share of emissions that stays in the atmosphere: atmospheric growth divided by total emissions. For the decade above it is 5.1 ÷ 11.2 ≈ 0.46. AR6 noted that land and ocean sinks together have taken up about 56 per cent of human carbon dioxide emissions over the last six decades, so the long-run airborne fraction is about 44 per cent. The fraction has stayed broadly constant because sinks have grown as emissions grew; AR6 projects that in scenarios with higher emissions the sinks take up a larger amount but a smaller fraction. Carbon dioxide removal by natural sinks is not the same as a lifetime. There is no single atmospheric lifetime for carbon dioxide: a pulse is drawn down partly within decades by the ocean surface and land, partly over centuries by the deep ocean, and partly only over many thousands of years by carbonate dissolution and weathering, so a substantial fraction remains for millennia. That long tail is why carbon dioxide warming depends on cumulative emissions. Isotopic and oxygen evidence shows that the rise is from fossil carbon. Fossil carbon is free of carbon-14 and depleted in carbon-13, and atmospheric carbon dioxide has shown both a declining ¹³C/¹²C ratio (the Suess effect) and, before nuclear testing complicated it, a dilution of ¹⁴C. Atmospheric oxygen is falling in step with combustion. The north–south gradient of carbon dioxide, higher in the industrial Northern Hemisphere, adds further support.
Methane is the second largest contributor to human-caused forcing. Its global mean concentration was about 1866 parts per billion (ppb) in 2019, against about 730 ppb before industrialisation, a rise of more than 150 per cent. Global methane budgets such as those of the Global Carbon Project put total sources at roughly 550–600 Tg of methane per year, of which roughly 60 per cent is anthropogenic. Natural sources are led by wetlands, with smaller contributions from freshwaters, termites, geological seeps and wild animals. Anthropogenic sources are led by agriculture and waste (enteric fermentation in ruminants, manure management, flooded rice paddies, landfills and wastewater) and by fossil fuel production and use (leaks and venting from oil, gas and coal). Biomass burning contributes in both categories. The main sink is oxidation by the hydroxyl radical (OH) in the troposphere, which removes about 90 per cent; the rest is lost in the stratosphere and by uptake in dry soils. Because OH is itself consumed by methane, more methane slows its own removal, so the perturbation lifetime (about 12 years) is longer than the total atmospheric lifetime (about 9 years). At steady state, sources equal burden divided by lifetime: with a burden near 5200 Tg and a lifetime of about 9 years, the implied annual sink is close to 570 Tg, consistent with the bottom-up budgets. Methane oxidation also produces tropospheric ozone and stratospheric water vapour, both of which add warming; these indirect effects are included in its GWP. Atmospheric methane growth paused in the early 2000s and resumed afterwards. The isotopic signal of the renewed rise has shifted towards lighter carbon, which points to a substantial biogenic contribution from wetlands, agriculture or waste, but the relative roles of these and of changes in OH are still debated. Satellites and top-down inversions increasingly identify individual large leaks, so-called super-emitters, in oil and gas fields.
Nitrous oxide reached about 332 ppb in 2019, against about 270 ppb before industrialisation. It has a long atmospheric lifetime, about 109 years in AR6, and its main sink is photolysis and reaction with excited oxygen atoms in the stratosphere. That reaction also produces nitric oxide, which destroys stratospheric ozone; since chlorofluorocarbons were phased out under the Montreal Protocol, nitrous oxide has become the dominant ozone-depleting substance emitted, yet it is not controlled under that treaty. Total sources are of the order of 17 Tg of nitrogen per year. Natural sources are soils under natural vegetation and the ocean. The main anthropogenic source is agriculture: microbial nitrification and denitrification in soils supplied with synthetic fertiliser nitrogen, manure and crop residues, plus manure management. Industry (adipic and nitric acid production), fossil fuel combustion, biomass burning and wastewater add the rest. Emissions are episodic and spatially patchy: a large share of a field's annual emission can occur in a few days after fertiliser, rain or thaw, which makes measurement hard. The IPCC Guidelines for National Greenhouse Gas Inventories estimate direct soil emissions at Tier 1 with an emission factor, EF1. In the 2006 Guidelines EF1 is 0.01 kg N2O–N per kg of nitrogen applied, that is, 1 per cent. The 2019 Refinement keeps an aggregate default of 1 per cent but gives disaggregated values, higher for synthetic fertiliser in wet climates (1.6 per cent) and lower in dry climates (0.5 per cent). To convert N2O–N to N2O, multiply by 44/28. Indirect emissions from leached and volatilised nitrogen are calculated separately. Because the multiplier is large (GWP100 of 273 in AR6), small changes in nitrogen use efficiency matter for a farm's footprint.
An aerosol is a suspension of fine solid or liquid particles in air. Natural aerosols include sea salt, mineral dust, volcanic sulphate and secondary organic particles formed from plant emissions. Anthropogenic aerosols come mainly from fossil fuel and biomass burning: sulphate formed from sulphur dioxide, nitrate, black carbon (soot), organic carbon and ammonium formed from agricultural ammonia. Most tropospheric aerosols are removed by rain within days to a couple of weeks, so their concentrations are patchy and closely track current emissions, unlike the well-mixed greenhouse gases. Aerosols affect climate in two main ways. Aerosol–radiation interactions (the 'direct effect') are scattering and absorption of sunlight. Scattering aerosols such as sulphate reflect sunlight and cool; absorbing aerosols such as black carbon warm the atmospheric layer they are in and darken snow and ice when deposited. Absorption can also alter cloud cover by heating the layer, sometimes called the semi-direct effect. Aerosol–cloud interactions (the 'indirect effects') arise because aerosols act as cloud condensation nuclei. For the same liquid water, more nuclei give more and smaller droplets and a brighter cloud; this is the Twomey effect. Smaller droplets may also suppress drizzle and lengthen cloud lifetime (the Albrecht or lifetime effect), though observations show cloud responses that partly offset this. AR6 assessed aerosol ERF from 1750 to 2014 at about −1.3 W m⁻² (very likely range −2.0 to −0.6), and about −1.1 W m⁻² for 2019 as emissions in some regions fell; aerosol–cloud interactions provide the larger part. This is the largest uncertainty in total human forcing, and through it in the observational estimate of climate sensitivity.
Large explosive eruptions inject sulphur dioxide into the stratosphere, where it forms sulphate aerosol that can persist for one to two years because there is no rain to remove it. The global cooling that follows is a natural experiment in forcing and response: after the 1991 Mount Pinatubo eruption global mean surface temperature fell by roughly half a degree for a year or two. Tropospheric pollution, by contrast, washes out in days, so its cooling vanishes almost immediately when emissions stop. This asymmetry has practical meaning. Anthropogenic aerosols have masked part of greenhouse warming. AR6 assessed that well-mixed greenhouse gases contributed 1.0 to 2.0 °C of warming by 2010–2019 relative to 1850–1900, while other human drivers, mainly aerosols, contributed a cooling of 0.0 to 0.8 °C. Clean-air policies that cut sulphur dioxide improve health and reduce acid rain, but they also remove part of this mask, adding some near-term warming. Cuts in methane and black carbon alongside sulphur cuts can offset it. Regional effects differ from the global mean: aerosols have influenced monsoon rainfall in South Asia and the Sahel and alter the distribution of heating between hemispheres. Deliberately adding aerosol to the stratosphere to reflect sunlight, called stratospheric aerosol injection or solar radiation modification, is discussed in AR6 as a possible response option. It would not address ocean acidification, could change regional rainfall, and would carry a risk of rapid 'termination' warming if stopped abruptly while greenhouse gas concentrations stayed high. It is governed by no binding international framework and remains highly controversial.
A feedback is a process that changes in response to warming and in turn changes the radiation balance. The standard linear framework writes the top-of-atmosphere imbalance as N = F + λΔT, where F is the effective radiative forcing, ΔT the change in global mean surface temperature, and λ the net feedback parameter in W m⁻² K⁻¹, negative for a stable climate. At equilibrium N = 0, so ΔT = −F/λ. The equilibrium climate sensitivity (ECS) is the equilibrium warming for a doubling of carbon dioxide: ECS = −F₂ₓ/λ. The basic stabilising response is the Planck feedback: a warmer planet radiates more. For a black body at 255 K the derivative of σT⁴ is 4σT³, about 3.8 W m⁻² K⁻¹; AR6 assessed the Planck feedback at about −3.22 W m⁻² K⁻¹ once its geographic pattern is considered. With the Planck response alone, a doubling of carbon dioxide would warm the climate by only about 1.2 °C. The other feedbacks change this. The water vapour feedback is strongly positive: by the Clausius–Clapeyron relation, saturation vapour pressure rises by about 7 per cent per kelvin, and relative humidity stays roughly constant, so a warmer atmosphere holds more of the most important greenhouse gas. The lapse-rate feedback is negative in the tropics, where the upper troposphere warms more than the surface; because the two are physically linked they are often combined. The surface-albedo feedback is positive, as snow and sea ice retreat and expose darker surfaces. The cloud feedback is the most uncertain; AR6 assessed it as positive, mainly from low clouds over subtropical oceans and rising high cloud tops. AR6 central values are about +1.30 W m⁻² K⁻¹ for water vapour plus lapse rate, +0.35 for surface albedo and +0.42 for clouds, giving a net feedback near −1.16 W m⁻² K⁻¹.
For four decades the likely range of ECS was 1.5–4.5 °C, from the Charney report of 1979 through AR5. AR6 narrowed it to a likely range of 2.5–4.0 °C, a very likely range of 2.0–5.0 °C, and a best estimate of 3.0 °C. The narrowing came not from one line of evidence but from combining several: process understanding of feedbacks, especially clouds; the historical record of warming and ocean heat uptake; and palaeoclimate states such as the Last Glacial Maximum, about 21,000 years ago, when the globe was about 5–7 °C colder, and warm periods such as the mid-Pliocene. AR6 did not base its range directly on the spread of climate models; some newer models have ECS above 5 °C, outside the assessed very likely range. The transient climate response (TCR) is the warming at the time carbon dioxide doubles in a run where it rises by 1 per cent per year, about 70 years. It is smaller than ECS because the ocean is still taking up heat. AR6 gives a best estimate of 1.8 °C with a likely range of 1.4–2.2 °C. TCR is the more relevant number for warming over this century; ECS is the more relevant number for long-term commitment. Feedbacks are not fixed. The 'pattern effect' describes how the geographic pattern of sea surface warming changes the strength of feedbacks; the observed pattern of recent decades, with relatively little warming in the eastern tropical Pacific, has given apparently weaker feedbacks than are expected later. This is why estimates of ECS from the historical record alone tend to come out low. Separately, Earth system sensitivity, which includes slow ice-sheet and vegetation changes over millennia, is larger than ECS.
Global surface temperature is estimated by several independent groups who combine land station air temperature with sea surface temperature from ships and buoys, using different methods for gaps and bias corrections; their results agree closely. AR6 assessed global surface temperature in 2011–2020 at 1.09 °C (very likely 0.95–1.20) above 1850–1900, the period used as an approximation for pre-industrial. Land warmed more, about 1.59 °C, than the ocean surface, about 0.88 °C, because land has less heat capacity and less evaporative cooling. Each of the last four decades was successively warmer than any decade before it since 1850. AR6 assessed that temperatures during the most recent decade exceed those of any multi-century period since about 125,000 years ago, the last interglacial. The Arctic has warmed at more than twice the global rate, a pattern called Arctic amplification, driven largely by the loss of snow and sea ice. Satellite and radiosonde records show warming of the troposphere and cooling of the stratosphere. The water cycle has intensified. A warmer atmosphere holds more moisture, so heavy precipitation draws on more water: AR6 found with high confidence that the frequency and intensity of heavy precipitation events have increased since the 1950s over most land areas with good data. Global mean precipitation rises more slowly than water vapour, at about 1–3 per cent per kelvin, because it is limited by how fast the atmosphere can radiate away the latent heat released. The result is a tendency for wet regions and seasons to get wetter and for the contrast between wet and dry to sharpen, though the pattern over land is strongly modified by circulation. Hot extremes, including heatwaves, have become more frequent and more intense across most land regions since the 1950s, and cold extremes less so. Agricultural and ecological droughts have increased in some regions.
Global mean sea level rose by about 0.20 m between 1901 and 2018. The rate has accelerated: about 1.3 mm per year over 1901–1971, 1.9 mm per year over 1971–2006 and 3.7 mm per year over 2006–2018. Sea level rises because ocean water expands as it warms (thermal expansion) and because land ice melts and flows into the sea; the melting of floating sea ice does not raise sea level significantly. For 1971–2018 AR6 attributed about 50 per cent of the rise to thermal expansion, 22 per cent to glacier loss, 20 per cent to ice-sheet loss and 8 per cent to changes in land water storage. Ice-sheet loss has grown fastest: the rate of loss from Greenland and Antarctica together increased by a factor of four between 1992–1999 and 2010–2019, and ice sheets became the main contributor to sea level rise in the most recent period. Measurements come from tide gauges since the eighteenth century and from satellite altimetry since 1993; gravity satellites measure ice and water mass change. Relative sea level at any coast also depends on vertical land motion. Subsiding deltas, where groundwater or hydrocarbon extraction and sediment starvation lower the land, can see relative rise several times the global mean, while land still rebounding from the weight of past ice sheets can see falling relative sea level. The cryosphere shows some of the clearest signals. Arctic sea ice area has declined in all months, most in late summer; AR6 assessed that late-summer Arctic sea ice area in 2011–2020 was the lowest for at least 1000 years. Antarctic sea ice shows no significant overall trend over the satellite era to 2020, with large regional contrasts, followed by record lows after 2016. Glaciers worldwide have retreated in a way that AR6 judged unprecedented in at least the last 2000 years. Northern Hemisphere spring snow cover has decreased since 1950, and permafrost temperatures have increased to record levels. Ice core, tree ring, coral and sediment proxies give the long context against which these changes are judged.
The ocean is where most of the extra energy goes, so ocean heat content is the most robust indicator of a warming planet. It is measured by ship profiles and, since the mid-2000s, by the Argo array of several thousand profiling floats that sample the upper 2000 m. AR6 assessed it as virtually certain that the global upper ocean (0–700 m) has warmed since the 1970s and extremely likely that human influence is the main driver. The ocean warms from the top down; heat reaches the deep ocean mainly where water sinks in the North Atlantic and around Antarctica. Because so much heat has already entered the ocean and will keep entering it, ocean warming and the sea level rise from thermal expansion will continue for centuries even if emissions stop. Marine heatwaves, periods of extreme high sea surface temperature lasting days to months, have roughly doubled in frequency since the 1980s according to AR6, and they have caused mass coral bleaching, kelp forest loss and fishery disruption. Warmer water holds less dissolved oxygen, and warming of the surface increases stratification, which reduces the mixing of oxygen-rich surface water into the interior. Together these cause ocean deoxygenation. Oxygen content has declined in many regions since the mid-twentieth century, with the global loss in the upper ocean estimated at a few per cent, and oxygen minimum zones have expanded. Coastal hypoxic 'dead zones' are a separate but compounding problem, driven mainly by nutrient run-off from agriculture and sewage that feeds algal blooms whose decay consumes oxygen. Deoxygenation compresses the habitat of large fish such as tuna into shallower water, alters nitrogen cycling and can increase marine production of nitrous oxide.
Ocean acidification is the fall in seawater pH caused by uptake of carbon dioxide. When carbon dioxide dissolves it forms carbonic acid, which releases hydrogen ions; these combine with carbonate ions to form bicarbonate. The result is a lower pH and a lower carbonate-ion concentration. Open-ocean surface pH has fallen by about 0.1 units since pre-industrial times, from about 8.2 to about 8.1, with AR6 assessing it as virtually certain that this decline is occurring and stating with high confidence that current values are unusual for at least 26,000 years. Because pH is a logarithmic scale, a fall of 0.1 means hydrogen-ion concentration has risen by a factor of 10^0.1, about 1.26, or roughly 26–30 per cent. The ocean is not becoming acidic in the strict sense: its pH remains above 7. 'Acidification' describes the direction of change. Organisms that build calcium carbonate shells or skeletons are most exposed. The saturation state, Ω, compares the product of calcium and carbonate ion concentrations in seawater with the equilibrium solubility product of a mineral. Where Ω is above 1, the mineral tends not to dissolve; below 1, unprotected shells and skeletons tend to dissolve. Aragonite, the form used by corals and many molluscs and pteropods, is more soluble than calcite, so it becomes undersaturated first, and cold high-latitude and upwelling waters are already near or below aragonite saturation in places. Laboratory and field studies show reduced calcification and larval survival in many species, with wide variation between species. Shellfish hatcheries on upwelling coasts have already had to manage intake water chemistry. Acidification, warming and deoxygenation act together, a combination sometimes called the 'deadly trio'. Warm-water coral reefs are exposed to all three plus local pollution. The IPCC Special Report on Global Warming of 1.5 °C projected that warm-water coral reefs would decline by a further 70–90 per cent at 1.5 °C of global warming and by more than 99 per cent at 2 °C. Unlike warming, acidification is a direct chemical consequence of carbon dioxide concentration and would not be prevented by any measure that reflects sunlight.
Detection is the process of showing that climate has changed in some defined statistical sense, without giving a reason. Attribution is the process of evaluating the relative contributions of multiple causal factors to a detected change, with an assigned confidence. The two depend on estimates of internal variability, from long control runs of climate models and from observations, to judge whether a change could have arisen by chance. The standard method is optimal fingerprinting. Climate models produce the expected space–time pattern of response, the 'fingerprint', to each forcing: greenhouse gases, aerosols, solar and volcanic. The observed change is regressed onto these fingerprints, and the regression scaling factors tell whether each signal is present (a factor significantly above zero) and whether its size is consistent with the models (a factor consistent with one). Fingerprints separate causes that a simple global average cannot. A brighter Sun would warm the whole atmosphere, including the stratosphere; greenhouse gases warm the troposphere and cool the stratosphere, which is what is observed. Increased greenhouse gases also produce more warming at night than by day and in winter than in summer over many land regions. The confidence of these statements has risen in each assessment. AR6 stated that it is unequivocal that human influence has warmed the atmosphere, ocean and land. It assessed the human-caused increase in global surface temperature from 1850–1900 to 2010–2019 as about 1.07 °C (likely 0.8–1.3 °C), essentially all of the observed warming, with natural drivers contributing only −0.1 to +0.1 °C. Human influence is very likely the main driver of the global retreat of glaciers since the 1990s and the decrease in Arctic sea ice area, and of global mean sea level rise since at least 1971.
Event attribution asks how human-caused climate change has altered the likelihood or intensity of a particular kind of extreme event, such as a heatwave, flood-producing rainfall or drought. The usual risk-based approach compares the probability of an event at least as extreme as the one observed in today's climate, P1, with its probability in a counterfactual climate without human influence, P0. The probability ratio (or risk ratio) is PR = P1/P0. The fraction of attributable risk is FAR = 1 − P0/P1, the share of the present risk that is due to climate change. Equivalently, an intensity-based analysis asks how much hotter or wetter an event of the same rarity is now than it would have been. Methods combine statistical analysis of long observational records with large ensembles of climate model runs, and require model evaluation to check that the models simulate the relevant kind of event realistically. Results are stated for a defined event class, region and season, not for 'the storm' itself. Confidence is generally highest for heat extremes, because the thermodynamic link to warming is direct and models simulate them well; medium for heavy precipitation, which rises with atmospheric moisture but depends on circulation; and lowest for events such as droughts in some regions, tropical cyclone frequency or tornadoes, where data are short or models struggle. AR6 stated that some recent hot extremes observed in the past decade would have been extremely unlikely to occur without human influence. A complementary 'storyline' approach takes the meteorology of a particular event as given and asks how the warmer, moister background changed its outcome. It answers a different question from the probabilistic approach and is useful for planning. Attribution results feed into loss-and-damage discussions, adaptation planning, litigation and financial disclosure, so the precision of their wording matters: an event being made 'five times more likely' is a statement about P1/P0, not about the event having been caused solely by climate change.
One of the most policy-relevant findings of climate science is that global warming from carbon dioxide is almost linearly proportional to cumulative carbon dioxide emissions. The constant of proportionality is the transient climate response to cumulative emissions of carbon dioxide (TCRE). AR6 gives a best estimate of 1.65 °C per 1000 PgC, likely range 1.0–2.3 °C per 1000 PgC; in units of carbon dioxide this is 0.45 °C per 1000 GtCO2 (likely 0.27–0.63). The near-linearity arises because two effects roughly cancel: the forcing per extra tonne falls as carbon dioxide rises (the logarithm), while the share of each tonne staying in the air rises as sinks saturate and the ocean takes up heat more slowly. A remaining carbon budget is the cumulative net carbon dioxide that can still be emitted while keeping warming below a limit with a stated probability. It is calculated from the warming to date, TCRE, the projected warming from non-carbon dioxide forcers, the zero emissions commitment, and adjustments for Earth system feedbacks such as permafrost thaw. AR6 estimated historical cumulative emissions over 1850–2019 at about 2390 GtCO2 (±240). From the start of 2020 the remaining budget for a 50 per cent chance of limiting warming to 1.5 °C was 500 GtCO2 (400 GtCO2 for a 67 per cent chance), and for 2 °C it was 1350 GtCO2 at 50 per cent (1150 GtCO2 at 67 per cent). Global emissions of about 40 GtCO2 per year use up a 500 GtCO2 budget in about 12 years, which is why the 1.5 °C budget is now nearly exhausted. Uncertainties in non-carbon dioxide forcing, especially methane and aerosols, shift these numbers by about ±220 GtCO2. The linear relationship has a corollary: stopping warming caused by carbon dioxide requires net zero carbon dioxide emissions. The zero emissions commitment (ZEC), the further temperature change after net zero carbon dioxide is reached, is assessed as close to zero (likely −0.3 to +0.3 °C), because continued ocean heat uptake roughly balances continued drawdown of carbon dioxide. Temperature therefore stabilises at about the level reached, but it does not fall without net negative emissions.
AR6 defines a tipping point as a critical threshold beyond which a system reorganises, often abruptly and/or irreversibly. Irreversibility is judged on a stated time scale: a change is irreversible if recovery takes much longer than the time it took to cause it. Hysteresis means that reversing the forcing does not return the system along the same path; an ice sheet that has collapsed may need conditions much colder than those that triggered the collapse to regrow. Candidate tipping elements include the Greenland ice sheet, whose surface lowering exposes it to warmer air in a melt–elevation feedback; parts of the West Antarctic ice sheet, which rests on bedrock below sea level and deepening inland, making it vulnerable to marine ice sheet instability as the grounding line retreats; the Atlantic Meridional Overturning Circulation (AMOC), the system of northward warm surface flow and southward deep return flow that carries heat into the North Atlantic; the Amazon rainforest, where deforestation and drying could shift parts towards a degraded or savanna-like state; permafrost thaw, which releases carbon dioxide and methane; and warm-water coral reefs. AR6 assessed that AMOC is very likely to weaken over the twenty-first century but, with medium confidence, will not collapse abruptly before 2100. Thresholds for most elements are uncertain, and published estimates of them differ, so they are best treated as risks that rise with warming rather than as precise temperature lines. Some changes are already committed. AR6 stated that changes in the ocean, ice sheets and global sea level are irreversible for centuries to millennia. Sea level will continue to rise for centuries to millennia because of continuing deep-ocean warming and ice-sheet melt, and will remain elevated for thousands of years. Low-likelihood, high-impact outcomes such as ice-sheet collapse, abrupt ocean circulation change and much higher warming than the best estimate cannot be ruled out and are part of a risk assessment. For a risk manager the relevant fact is that the likelihood of such outcomes increases with every increment of warming.
The Intergovernmental Panel on Climate Change was established in 1988 by the World Meteorological Organization (WMO) and the United Nations Environment Programme (UNEP). It does not carry out its own research or monitor the climate. It assesses the published literature, peer-reviewed first and other sources where appropriate, and reports what is known, with what confidence, and where the gaps are. Its products are policy-relevant but not policy-prescriptive: they say what the consequences of options are, not which option governments should choose. Work is organised in three working groups and a task force. Working Group I assesses the physical science basis; Working Group II impacts, adaptation and vulnerability; Working Group III mitigation; and the Task Force on National Greenhouse Gas Inventories produces the methodology guidelines that countries use to report emissions, notably the 2006 Guidelines and the 2019 Refinement. Each assessment cycle produces a working group report from each group, a Synthesis Report, and usually several special reports. The sixth cycle produced the WGI report in 2021, WGII and WGIII in 2022, and the Synthesis Report in 2023, plus special reports on 1.5 °C warming, climate change and land, and the ocean and cryosphere. The seventh cycle is under way. Authors are nominated by governments and observer organisations and selected by the bureau for expertise and balance of region and gender. Drafts pass through successive expert and government reviews, and every comment receives a documented response. The Summary for Policymakers (SPM) is approved line by line by governments in a plenary session together with the authors; governments may ask for wording to be clearer, but the text must remain consistent with the underlying report, which is accepted rather than approved line by line. This process gives the SPM unusual authority: every government has signed off on its wording. The IPCC shared the Nobel Peace Prize in 2007.
IPCC authors use two calibrated scales, defined in a guidance note for the AR5 authors and used again in AR6. Confidence is a qualitative judgement of the validity of a finding, based on the type, amount, quality and consistency of evidence (summarised as limited, medium or robust) and the degree of agreement (low, medium or high). It is expressed in five levels: very low, low, medium, high and very high. Robust evidence with high agreement gives the highest confidence. Likelihood is a quantified probability, used when uncertainty can be expressed statistically or by expert judgement. The terms are: virtually certain, 99–100 per cent probability; extremely likely, 95–100 per cent; very likely, 90–100 per cent; likely, 66–100 per cent; about as likely as not, 33–66 per cent; unlikely, 0–33 per cent; very unlikely, 0–10 per cent; extremely unlikely, 0–5 per cent; and exceptionally unlikely, 0–1 per cent. 'More likely than not' (above 50 per cent) is also used. In the reports these terms are set in italics so the reader knows they carry the calibrated meaning. A 'likely range' is therefore a range with at least a 66 per cent chance of containing the true value; a 'very likely range' at least 90 per cent. The two scales are complementary, and a finding can carry both, such as a likely range assessed with high confidence. Practitioners misread the language in predictable ways. 'Likely' is not 'probably true and therefore certain'; there is up to a one-in-three chance the value lies outside a likely range. 'Low confidence' does not mean that something will not happen; it means evidence or agreement is too weak to judge. AR6 also uses 'unequivocal' for its strongest statement of fact about observed human-caused warming. In financial disclosure, risk scenarios and policy briefs, quoting IPCC language with its calibrated meaning, and not converting 'medium confidence' into 'proven', is a professional duty.
A student uses this subject to master the numerical core that entrance and postgraduate papers test: computing effective temperature from the solar constant and albedo, converting ppm to gigatonnes of carbon, converting nitrous oxide–nitrogen to nitrous oxide and then to CO2-equivalent with a named GWP set, applying the logarithmic forcing formula, and turning a feedback parameter into a sensitivity. The student should also learn to quote IPCC numbers with their assessment report and their likelihood range, because examiners set distractors from older reports (AR4 and AR5 GWPs, the 1.5–4.5 °C ECS range) and from confusions between confidence and likelihood.
A five-year practitioner, such as a carbon or LCA analyst, an agri-food sustainability officer or a climate-risk analyst in a bank, uses the subject daily to choose and disclose GWP sets, explain to managers why methane cuts give a fast temperature benefit, convert inventory activity data with IPCC emission factors, and translate IPCC findings into physical-risk narratives without overstating them. Typical decisions are whether a client's target is consistent with a remaining carbon budget, how to word an attribution finding in a disclosure, and whether to report GWP20 or GWP* alongside GWP100 for a livestock or gas portfolio.
A senior professional with 25 years' experience, such as a head of sustainability, a chief risk officer's climate lead or a senior regulator, uses the subject to set policy and challenge advice. That means judging whether an organisation's net-zero claim relies on offsets that do not match the permanence of fossil carbon, recognising when a consultant's scenario ignores tipping-point tail risks or the aerosol unmasking effect of clean-air rules, and making sure board papers use IPCC calibrated language. Such a person also explains to non-scientists why a stable methane herd and a growing fossil fleet are not equivalent, while resisting the use of GWP* to excuse high historical emissions.
A researcher with a PhD and 25 years' experience, such as a principal scientist, professor or IPCC author, works at the edges of this subject: constraining cloud feedbacks and the pattern effect, reducing aerosol forcing uncertainty, closing the methane budget with isotopes and satellite inversions, quantifying permafrost and wetland feedbacks for carbon budgets, refining extreme-event attribution for data-poor regions of Africa and South Asia, and assessing tipping thresholds. Such a researcher also designs the assessments themselves, writes the calibrated statements governments will approve, and supervises students across the physical, agricultural and economic sides of climate science.
This manual gives AR6 values; newer annual updates of the global carbon and methane budgets, the remaining carbon budget and observed temperature (including the record warmth of recent years) will differ in detail and should be read from the current Global Carbon Budget and WMO State of the Climate reports. It does not cover climate model construction, scenarios (SSPs and RCPs) or regional projections, which belong to the climate-models-scenarios subject, nor impacts on crops and adaptation, which have their own subjects. Exact GTP values, the full list of halocarbon GWPs, regional sea level projections and quantitative threshold temperatures for individual tipping elements are left out because they vary between studies and assessments. Radiative transfer is treated without equations beyond the Stefan–Boltzmann law and the simplified logarithmic CO2 formula. Ozone-layer science is mentioned only where it touches climate.
Original practice questions written to the syllabus and pattern of the exams named in each question's examTags. They are not past papers and do not reproduce any board's questions.
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