Science Wiki

The carbon cycle is the biogeochemical cycle by which carbon is exchanged, transformed, and conserved among the biosphere, pedosphere, lithosphere, hydrosphere, cryosphere, and atmosphere. Within Earth system science, the carbon cycle is both a principal control on planetary climate and the chemical framework for all known cellular life.

Carbon's electronic configuration (1_s_²2_s_²2_p_²) allows tetravalent covalent bonding through sp³, sp², and sp orbital hybridization, permitting the assembly of stable macromolecules including proteins, nucleic acids, lipids, and carbohydrates. In the environment, carbon occurs in oxidation states ranging from −4 in fully reduced methane (CH₄) to +4 in fully oxidized carbon dioxide (CO₂) and carbonate (CO₃²⁻). In the atmosphere, trace concentrations of CO₂ and CH₄ absorb outgoing longwave infrared radiation, contributing to the greenhouse effect. The global carbon cycle is conventionally divided into two coupled domains distinguished by reservoir residence time and flux magnitude: the fast (biogeochemical) cycle and the slow (geochemical) cycle.

The distribution of carbon on Earth is overwhelmingly skewed toward lithospheric reservoirs, which exchange mass slowly with the active surface pools (the ocean, atmosphere, and terrestrial biosphere). Box models of the carbon cycle express reservoir stocks in petagrams of carbon (PgC; 1 PgC = 10¹⁵ g = 1 gigatonne of carbon, GtC) and fluxes in PgC per year (PgC/yr).

Approximate sizes of the major carbon reservoirs

The fast carbon cycle operates on diurnal, seasonal, and centennial timescales through the biological oxidation and reduction of carbon among the atmosphere, terrestrial ecosystems, and the surface ocean.

Photoautotrophic organisms, chiefly vascular plants, cyanobacteria, and eukaryotic algae, drive the primary influx of carbon into the biosphere. In oxygenic photosynthesis, the light-dependent reactions in the thylakoid membranes of the chloroplast split water to generate adenosine triphosphate (ATP) and reduced nicotinamide adenine dinucleotide phosphate (NADPH). These cofactors power the Calvin cycle in the stroma, where the enzyme RuBisCO (ribulose-1,5-bisphosphate carboxylase/oxygenase) adds CO₂ to ribulose-1,5-bisphosphate, yielding two molecules of 3-phosphoglycerate. ATP and NADPH then reduce 3-phosphoglycerate to triose phosphates, the precursors of hexose sugars. The overall reaction is:

Most terrestrial plant biomass is produced by C₃ carbon fixation. Lineages adapted to high temperature, aridity, and low atmospheric CO₂ use C₄ fixation or crassulacean acid metabolism (CAM). C₄ plants, such as many tropical grasses, use phosphoenolpyruvate carboxylase to concentrate CO₂ in bundle-sheath cells, suppressing photorespiration. CAM plants separate CO₂ uptake (at night) from the Calvin cycle (by day) in time. Globally, terrestrial photosynthesis removes approximately 120 PgC/yr from the atmosphere, a flux termed gross primary production (GPP).

With approximately 38,000 PgC of dissolved inorganic carbon (DIC), the ocean is the largest carbon reservoir that exchanges actively with the atmosphere, holding roughly 40 times as much carbon as the present-day atmosphere.

Unlike N₂ and O₂, which dissolve without reacting with water, CO₂ reacts chemically after crossing the air–sea interface. Under Henry's law, CO₂ dissolves in proportion to its partial pressure:

Dissolved CO₂ hydrates to form carbonic acid, which dissociates into bicarbonate and carbonate ions:

Here [H₂CO₃^*] denotes the combined concentration of dissolved CO₂ and true carbonic acid. Total dissolved inorganic carbon is the sum of the three inorganic species:

At a typical surface seawater pH of ~8.1, DIC consists of approximately 90% bicarbonate, 9% carbonate, and about 1% dissolved CO₂. When additional CO₂ enters the ocean, the released protons are largely neutralized by carbonate ions, giving the net buffering reaction:

This buffering gives the ocean its large capacity to store carbon. However, uptake of CO₂ consumes carbonate ions while leaving total alkalinity (TA) unchanged:

As carbonate is depleted, the ocean's chemical capacity to absorb further CO₂ declines. This is quantified by the dimensionless Revelle factor (RF), the ratio of the fractional change in seawater _p_CO₂ to the fractional change in DIC:

RF ranges from about 9 in warm tropical waters to about 15 in cold polar waters. With RF ≈ 10, a 10% increase in atmospheric _p_CO₂ produces only about a 1% increase in equilibrium surface-ocean DIC.

The solubility pump is a physicochemical process arising from the coupling of temperature-dependent gas solubility with the ocean's overturning circulation. Because the dissolution of CO₂ is exothermic, the solubility constant K₀ increases substantially as water cools.

As surface waters move to high latitudes in the North Atlantic and around Antarctica, they lose heat to the atmosphere and absorb additional CO₂. Cooling, together with brine rejection during sea-ice formation, increases seawater density, and these waters sink to form North Atlantic Deep Water (NADW) and Antarctic bottom water (AABW). This downwelling isolates dissolved inorganic carbon in the deep ocean for roughly 500 to 1,500 years. Deep waters return to the surface mainly through wind-driven upwelling in the Southern Ocean, with equatorial upwelling of shallower waters, where warming reduces solubility and CO₂ outgasses to the atmosphere.

Biological activity maintains a vertical gradient in DIC, keeping surface-ocean DIC roughly 10 to 15% lower than deep-ocean DIC, through three mechanisms: _F_(z) = _F_(z₀) · (z / z₀)^−_b_ where the attenuation exponent b is typically about 0.86. On the order of 1% of surface primary production reaches the deep seafloor, and only ~0.1 to 0.2 PgC/yr (less than 0.5% of marine NPP) is ultimately buried, mostly in continental-margin sediments.

• Carbonate counter-pump. Calcifying organisms, including coccolithophores, foraminifera, pteropods, and scleractinian corals, precipitate calcium carbonate as calcite or aragonite:

Ca²⁺ + 2HCO₃⁻ → CaCO₃(s) + CO₂(aq) + H₂O Precipitating one mole of CaCO₃ lowers DIC by one mole but lowers total alkalinity by two equivalents, shifting the carbonate equilibrium toward dissolved CO₂. Calcification therefore raises surface-water _p_CO₂, hence the term "counter-pump". Dense carbonate minerals also act as ballast, increasing the sinking speed of organic aggregates. Below the saturation horizon, sinking carbonate particles encounter undersaturated waters and begin to dissolve; dissolution increases markedly below the lysocline, and little carbonate accumulates below the carbonate compensation depth (CCD). This dissolution returns alkalinity to the deep ocean.

• Microbial carbon pump. Heterotrophic bacteria and archaea, together with viral lysis (the viral shunt), convert labile dissolved organic carbon into recalcitrant dissolved organic carbon (RDOC) that resists further degradation. RDOC makes up most of the ~700 PgC ocean DOC pool, which has a mean radiocarbon age in the deep ocean of about 4,000 to 6,000 years.

Operating over 10⁵ to 10⁸ years, the slow carbon cycle governs the exchange of carbon between the solid Earth (crust and mantle) and the surface reservoirs, and has helped maintain Earth's long-term habitability.

The principal long-term sink for atmospheric CO₂ is the chemical weathering of calcium and magnesium silicate minerals on the continents, commonly represented by wollastonite (CaSiO₃). Weathering of carbonate rocks, by contrast, has no net effect on atmospheric CO₂ over long timescales, because the CO₂ consumed during weathering is released again when the carbonate reprecipitates in the ocean.

In the carbonate–silicate cycle, atmospheric CO₂ dissolves in rainwater and soil water to form carbonic acid, which attacks silicate minerals through silicate weathering:

Rivers carry the dissolved products to the ocean, where calcifying organisms precipitate CaCO₃ and silicifying organisms such as diatoms and radiolarians precipitate biogenic opal (SiO₂):

The sum of these two reactions is the Urey reaction, which removes one mole of CO₂ from the ocean–atmosphere system per mole of calcium silicate weathered and stores it in marine carbonate sediments:

Plate tectonics carries carbonate-bearing sediments and altered oceanic crust into subduction zones. There, high temperatures (typically several hundred degrees Celsius) and pressures drive metamorphic decarbonation, the reverse of the Urey reaction, releasing CO₂. Some subducted carbon is returned to the deep mantle. Metamorphic CO₂, together with mantle carbon degassed at mid-ocean ridges, volcanic arcs, and hotspots, returns to the atmosphere at a global rate of about 0.1 PgC/yr.

This cycle acts as a negative-feedback thermostat, sometimes called the Walker feedback. If volcanic outgassing or solar luminosity increases, the climate warms. Higher temperatures accelerate silicate dissolution kinetics and intensify the hydrological cycle, increasing precipitation and runoff. Faster weathering then draws down atmospheric CO₂ over roughly 10⁵ to 10⁶ years until a new balance between outgassing and weathering is reached.

A second geological pathway is the burial of organic matter (represented as CH₂O). When marine or lake detritus is rapidly buried in anoxic sediments, it escapes aerobic decomposition. Through diagenesis and thermal maturation over millions of years, most of this material is converted to insoluble kerogen in shales, while a smaller fraction forms petroleum, natural gas, and coal.

Each mole of organic carbon buried represents a mole of photosynthetically produced O₂ that was not consumed by respiration. Burial of organic carbon, together with burial of sedimentary pyrite, is therefore the main long-term source of free oxygen in Earth's atmosphere. The cycle closes when orogeny uplifts organic-rich rocks to the surface, where oxidative weathering of kerogen (petrogenic carbon oxidation) consumes O₂ and returns CO₂ to the atmosphere.

Carbon has two stable isotopes, ¹²C (98.93%) and ¹³C (1.07%), and a trace cosmogenic radioisotope, ¹⁴C (radiocarbon; half-life 5,730 ± 40 years), produced in the upper atmosphere when thermal neutrons react with ¹⁴N. Isotope ratios serve as tracers of carbon cycle processes on both modern and geological timescales. Stable isotope composition is expressed in δ¹³C notation, in parts per thousand (‰), relative to the Vienna Pee Dee Belemnite (VPDB) standard:

During photosynthesis, kinetic isotope fractionation discriminates against ¹³C. Molecules of ¹²CO₂ diffuse slightly faster through stomata (a fractionation of about 4.4‰). Molecules containing the lighter isotope also have a higher zero-point vibrational energy, so their bonds are more easily broken and they react faster at the active site of RuBisCO (a fractionation of about 27–30‰). Typical values are:

• Atmosphere and carbonates: pre-industrial atmospheric CO₂ had a δ¹³C of about −6.5‰ (present-day ≈ −8.6‰). Marine carbonates precipitated near equilibrium with seawater bicarbonate have δ¹³C ≈ 0‰.
• Organic matter: C₃ plants produce biomass with δ¹³C of about −25 to −28‰, and C₄ plants about −12 to −14‰. Biogenic methane is strongly depleted in ¹³C, with δ¹³C of about −50 to −110‰.

In paleoclimatology, abrupt negative carbon isotope excursions (CIEs) recorded in marine carbonates indicate large releases of ¹³C-depleted carbon into the ocean–atmosphere system, as during the Paleocene–Eocene Thermal Maximum (PETM, ~56 million years ago).

Since the Industrial Revolution, burning fossil fuels and producing cement from limestone have transferred lithospheric carbon to the atmosphere roughly 100 times faster than the natural rate of volcanic outgassing.

By conservation of mass, emissions from fossil fuels and industry (E_FOS, ~10 PgC/yr in the early 2020s) and from land-use change such as deforestation (E_LUC, ~1.1 PgC/yr) are partitioned among growth in the atmosphere (G_ATM), the ocean sink (S_OCEAN), and the land sink (S_LAND), with a residual budget imbalance (B_IM) reflecting uncertainties in the estimates:

The airborne fraction, G_ATM / (E_FOS + E_LUC), has remained near 45%, with the land biosphere absorbing about 30% and the ocean about 25% of annual emissions. Atmospheric CO₂ has risen from a pre-industrial Holocene level of ~278 parts per million (ppm) to more than 420 ppm, and atmospheric CH₄ from ~720 parts per billion (ppb) to more than 1,900 ppb.

Several independent lines of evidence show that the rise in atmospheric CO₂ comes from fossil fuels, including the Suess effect:

1. ¹³C dilution. Fossil fuels derive from ancient photosynthetic organic matter and are depleted in ¹³C (δ¹³C typically −20 to −30‰ for coal and oil, and lower for natural gas). Their combustion has shifted atmospheric δ¹³C from about −6.5‰ in 1750 to about −8.6‰ today.
2. ¹⁴C dilution. Fossil fuels have been isolated for millions of years, far longer than the half-life of ¹⁴C, so they contain no radiocarbon. Their combustion dilutes the atmospheric ¹⁴C/¹²C ratio.
3. Oxygen decline. Atmospheric O₂ is falling at the rate expected from fossil-fuel combustion.
4. Spatial gradient. CO₂ concentrations are higher in the Northern Hemisphere, where most fossil-fuel emissions occur.

In Earth system models, the future strength of the land and ocean carbon sinks depends on two opposing feedbacks:

• Carbon–concentration feedback (β > 0; a negative feedback on climate): higher atmospheric CO₂ increases the air–sea partial-pressure gradient that drives ocean uptake. On land, it stimulates C₃ photosynthesis through the CO₂ fertilization effect, increasing carboxylation rates, suppressing photorespiration, and improving water-use efficiency by allowing partial stomatal closure.
• Carbon–climate feedback (γ < 0; a positive feedback on climate): warming weakens both sinks. In the ocean, warmer water holds less CO₂ (lower K₀), and stronger thermal stratification reduces vertical mixing and deep-water ventilation. On land, warming accelerates heterotrophic respiration relative to photosynthesis, increases drought-induced tree mortality and wildfire emissions, and thaws permafrost, releasing previously frozen carbon as CO₂ and CH₄.

Because the ocean absorbs about a quarter of anthropogenic CO₂ emissions, the formation of carbonic acid increases the hydrogen ion concentration of seawater, causing ocean acidification. Mean surface-ocean pH has fallen from about 8.2 before the Industrial Revolution to about 8.1, corresponding to an increase of about 30% in [H⁺].

The added protons combine with carbonate ions to form bicarbonate, lowering the carbonate ion concentration and therefore the saturation state (Ω) of calcium carbonate minerals:

where _K_′_sp is the stoichiometric solubility product of aragonite or calcite; aragonite is about 50% more soluble than calcite. Seawater with Ω > 1 is supersaturated, favouring calcification, whereas seawater with Ω < 1 is undersaturated, and unprotected carbonate shells and skeletons tend to dissolve. Acidification lowers Ω_arag and Ω_calc in surface waters and causes the aragonite and calcite saturation horizons to shoal (move upward). Over longer timescales the carbonate compensation depth is also expected to shoal. These changes raise the energetic cost of calcification for calcifying plankton and benthic invertebrates.

• Biological pump
• Carbon sequestration
• Climate change
• Nitrogen cycle
• Permafrost carbon cycle