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Carbon cycling on Earth
Aside from helium, living and universal matter are principally composed of the same chemical elements: hydrogen, oxygen, carbon, and nitrogen.
In the order of abundance: C - O - H - N (living) vs H - O - C - N (universal).
In contrast, the Earth has an entirely different composition: iron - oxygen - silicon - magnesium.
Interstellar clouds are composed of gases and dust grains in a mass ratio of about 100:1. The most dominant molecular species is hydrogen, in a concentration of about 104 H2 per cm3. The dust grains, measuring less than 0.5 μ in diameter, are in all probability graphite, Fe-Mg silicates and iron. The grains might be covered by solid water, ammonia and methane. The solid grain particles may provide protection for certain atoms or molecules, and also because their crystalline order catalysis and epitaxis may proceed favoring formation of carbon-containing molecules.
On transit to Earth, all carbon must have been present in a reduced state due to the large excess of hydrogen.
Upon arrival on Earth, the reduced carbon pool became part of the differentiation process leading to: core, mantle, crust, atmosphere and hydrosphere. As a result, carbon distributed itself among the various compartments. New molecular carbon species developed in accordance with changing mineral equilibria, while other species survived the harsh treatment of the evolving Earth molecularly intact.
Depending on pressure and temperature Cn takes form of graphite, carbynes and diamond.
Within the graphite stability field, graphite may react with hydrogen: C + 2H2 -> CH4 or with water vapor: C + H2O -> CO + H2. The subsequent reaction of carbon monoxide with water vapor will yield CO2: CO + H2O -> CO2 + H2.
Starting from graphite, a series of hydrocarbons and oxy-hydroxy compounds, such as graphitic acid C8O2(OH)2.
In summary, graphite is a compound which can combine with both hydrogen and oxygen, and equilibrium should exist between these molecular bonding states: [C-C]:[C-OH]:[C=O]:[CO2]:[H2O]:[H2].
In the past, CO2 was ultimately derived from mantle and crust. The present annual net flux of carbon dioxide from depth to surface, estimated at 0.06X1015 g C. For coparison, more than 100 times of this amount is yearly discharged into the atmosphere due to fossil fuel combustion and deforestation. It is significant that magmatic CO2 principally goes into the ocean, whereas man contaminates the atmosphere. Estimates on the pre-industrial CO2 level in our atmosphere range from 560 to 610 x 1015 g C compared to 730 x 1015 g C at present. These numbers indicate that more than half of man-made CO2 has already been picked up by a common sink, whether ocean, biota, or sediment.
Why CO2 is a greenhouse gas
"An enhanced greenhouse effect from CO2 has been confirmed by multiple lines of empirical evidence. Satellite measurements of infrared spectra over the past 40 years observe less energy escaping to space at the wavelengths associated with CO2. Surface measurements find more downward infrared radiation warming the planet's surface. This provides a direct, empirical causal link between CO2 and global warming."
"The greenhouse effect occurs because greenhouse gases let sunlight (shortwave radiation) pass through the atmosphere. The earth absorbs sunlight, warms then reradiates heat (infrared or longwave radiation). The outgoing longwave radiation is absorbed by greenhouse gases in the atmosphere. This heats the atmosphere which in turn re-radiates longwave radiation in all directions. Some of it makes its way back to the surface of the earth. So with more carbon dioxide in the atmosphere, we expect to see less longwave radiation escaping to space at the wavelengths that carbon dioxide absorb. We also expect to see more infrared radiation returning back to Earth at these same wavelengths."
More "physical" explanation of greenhouse effect: CHAPTER 7. THE GREENHOUSE EFFECT (Harvard)
Carbon dioxide infrared spectrum
Coblentz Society, Inc., "Evaluated Infrared Reference Spectra" in NIST Chemistry WebBook, NIST Standard Reference Database Number 69, Eds. P.J. Linstrom and W.G. Mallard, National Institute of Standards and Technology, Gaithersburg MD, 20899, http://webbook.nist.gov, (retrieved October 10, 2015). Carbon Dioxide - Infrared Spectrum
- The band saturation effect: Saturation, Nonlinearity and Overlap in the Radiative Efficiencies of Greenhouse Gases
- The saturation band broadening (Doppler shifting, pressure broadening): Why some gases are greenhouse gases, but most aren't, and some are stronger than others (.pdf)
- Layer model: A Multilayer Atmosphere Model (American Chemical Society)
- Climate sensitivity: Climate Sensitivity (American Chemical Society)
For CO2, the energy intensity Iout in units W/m2 goes up proportionally to the log of CO2 concentration. The logarithmic dependence means that you get about the same Iout change in W/m2 from any doubling of the CO2 concentration. The radiative effect of going from 10 to 20 ppm pCO2 is about the same as going from 100 to 200 ppm or 1,000 to 2,000 ppm.
Current CO2 measurements
Lifetime of carbon dioxide in the atmosphere
The rate of global fossil fuel CO2 emission grew at ~1% per year from 1980 to 2000 and >3% per year in the period from 2000 to 2005. Current emissions of major non-CO2 greenhouse gases such as methane or nitrous oxide are significant for climate change in the next few decades or century, but these gases do not persist over time in the same way as carbon dioxide.
There are rival definitions of a lifetime of a lifetime for anthropogenic CO2. One is the average amount of time that individual carbon atoms spend in the atmosphere before they are removed, by uptake into the ocean or terrestrial biosphere. Another is the amount of time it takes until the CO2 concentration in the air recovers substantially toward its original concentration. The difference between the two definitions is that exchange of carbon between the atmosphere and other reservoirs affects the first definition, by removing specific CO2 molecules, but not the second because exchange does not result in net CO2 drawdown. The misinterpretation that has plagued the question of atmospheric lifetime of CO2 seems to arise from confusion of these two very different definitions.
The models presented in Archer D et al. (Atmospheric Lifetime of Fossil Fuel Carbon Dioxide. Annual Review of Earth and Planetary Sciences. Vol. 37: 117-134, May 2009) give a broad picture of the fate of fossil fuel CO2 released into the atmosphere. Equilibration with the ocean will absorb most of it on timescale of 2 to 20 centuries. Even if this equillibration were allowed to run to completion, a substantial fraction of the CO2 (10-40%) would remain in the atmosphere awaiting slower chemical reaction with CaCO3 and igneous rocks. The remaining CO2 is abundant enough to continue to have a significant impact on climate for thousands of years. The changes in climate amplify themselves somewhat by driving CO2 out of the warmer ocean. The CO2 invasion has acidified the ocean, the pH of which is largely restored by excess dissolution of CaCO3 from the sea floor and on land and, ultimately, by silicate weathering on land. The recovery of ocean pH restores the ocean's buffer capacity to absorb CO2, tending to pull CO2 toward lower concentrations over the next 10,000 years. The land biosphere has its greatest impact within the first few centuries, which is when CO2 peaks. Nowhere in these model results or in the published literature is there any reason to conclude that the effects of CO2 release will be substantially confined to just a few centuries. In contrast, generally accepted modern understanding of the global carbon cycle indicates that the climate effects of CO2 releases to the atmosphere will persist for tens, if not hundreds, of thousands of years into the future.
History of carbon dioxide in the atmosphere
Hadean to early Proterozoic
Shaw GH. Earth's atmosphere – Hadean to early Proterozoic. Chemie der Erde 68 (2008) 235-264 (.pdf)
The most ancient atmosphere cannot be considered in isolation from hydrosphere, because exchanges of materials between the two take place on time scales short compared to geologic times.
Prevailing view of last 25-50 years on the Earth's early atmosphere and its origin with a few variations is based on Holland's (1984) review and analysis. Please note that all accounts of atmospheric evolution are speculative for the lack of evidence. Most theories aim at reconstructions that would set reasonable limits on the range of the parameters that would be permissive to origin of life and subsequent evolution of life and atmosphere leading to present day atmosphere.
The Archean Eon is a geologic eon, 4 Ga to 2.5 Ga years ago, following the Hadean (4.5-4 Ga) and preceding the Proterozoic (2.5-0.5 Ga). During the Archean, the earth's crust and layers had just formed making the Earth much cooler than it was during the Hadean, allowing the formation of continents.
Earth's earliest atmosphere was the product of late accretionary and internal heat-driven processes acting on late-accreting material mixed with the outer (fairly hot) parts of the Earth. Fifty percent or more of Earth's volatiles may have been degassed probably before 4.2 Ga. Because the composition of the earliest atmosphere is relevant to the origin of prebiotic molecules and to the subsequent evolution of the atmosphere, it is important to assess the range of possible compositions. Although Holland (1984) suggested a very short-lived phase with strongly reducing conditions resulting from degassing of H2 and CO leading to formation of CH4 (very briefly) and an unspecified quantity of organic compounds condensed onto/into the ocean, he also proposed that degassing quickly shifted to more oxidized gases (H2O and CO2) following core separation. He also suggested some minimal production of CH4 through interaction of oceanic organics and silicates under hydrothermal conditions, but considered the effects of this on the atmosphere to be very minor. Because he did not specify the relative amounts of organic compounds and CO2 production, the amount of reducing capacity stored as organic compounds from the earliest phase of the atmosphere remains an open question. Most discussions of atmospheric composition during the Archean invoke an atmosphere relatively rich in CO2, with only minor amounts of reduced carbon species. This is a matter of some importance to the origin of life because of the difficulty of producing significant amounts of prebiotic molecules from such an atmosphere, at least without the addition of significant reducing capacity. It may be that the current state of knowledge allows for more reducing capacity at the Earth's surface than has been assumed since an early CO2-rich atmosphere became the accepted model. Early degassing processes include (at least):
- impact degassing;
- volcanic degassing involving equilibration between volatiles and magmas at temperatures of 1000K;
- hydrothermal degassing involving equilibration or reaction of volatiles and various silicate phases at temperatures of hundreds of degrees.
The late-accretionary Earth had a solid surface, because magma oceans (even for something as large as the moon-forming impact) cool very rapidly, solidifying almost instantly on a geologic time scale. Within a few million years of the moon-forming impact the surface temperature dropped enough to allow liquid (water) oceans, with the solid Earth mostly covered by ocean. Although the moon-forming impact may have blown away a large part of the previously formed atmosphere, some residue (especially water dissolved in the magma ocean) might have remained. Note that the water in such a magma ocean would be rapidly stripped back into the surface environment once a frozen rock surface formed and temperatures dropped enough for water vapor to condense from the atmosphere (again, a few million years at most). This is because surface magmatism from the convecting magma ocean would release volatiles at low ambient pressure at the surface after the massive (water) atmosphere condensed to liquid. The mass of this early ocean is unknowable, but could easily have been equal to the current ocean, if not greater. The ocean was probably global, with few if any continents and little continental crust in the strict sense. Emergent parts of the silicate surface were confined to island-arclike settings.
During Archean carbon dioxide became the dominant carbon species in the atmosphere, but its concentration during Archean time is unknown. Although there is very limited evidence to indicate precise levels of atmospheric CO2 throughout much of the Archean, analysis of weathering near the end of Archean suggests that CO2 levels by that time were lower than required to maintain moderate surface temperatures.
It would require about 0.2 atm equivalent of carbon dioxide to overcome the faint sun during the earliest times, and only slightly less with the increasing luminosity of the sun 1-2 billion years later. This concentration is thought to be unsustainable for geologic intervals, given the reactivity of CO2 with silicates. This problem continues to conflict with invoking CO2 as the agent to overcome the climate paradox.
For the less luminous sun at 4.5 Ga, equable climate could probably be maintained at 1500 ppm CH4 with CO2 perhaps 10-20 times above modern levels. The loss of methane from anoxic atmosphere is accompanied by formation of condensable compounds that rain out into the ocean, i.e., prebiotic organic compounds. At the high rate of methane destruction (2 x 1011 molecules/cm2 s), a surface carbon reservoir equivalent to 20 atmospheres of CH4 (approx. the carbon in the current near-surface carbon inventory) cycles once through the atmosphere in about 100-200 million years. Because the photochemical destruction of methane removes, on average, about one hydrogen from each methane molecule in making the condensed organic phases, the carbon must be cycled through the atmosphere at least 3-4 times before it is oxidized to CO2. Note, however, that loss of reduction capacity actually depends not on methane production but loss of hydrogen to space. Crude calculations indicate that methane could be maintained at effective greenhouse concentration in the early (at least prebiotic) atmosphere for up to billion years, provided hydrothermal processes regenerate the methane from the (rained-out) oceanic carbon pool at a sufficient rate. Although this is a little short of the length of time required to maintain enough of a reduced reservoir to get to the end of Archean, it is at least in the right range. After the first billion years biologic processes may have contributed to atmospheric methane production.
Past 300 Myr
Retallack GJ. A 300-million-year record of atmospheric carbon dioxide from fossil plant cuticles. Nature. 2001 May 17;411(6835):287-90.
Retallack (2002) uses as a proxy for global palaeotemperature the inverse relationship between temperature and oxygen isotopic value (δ18O) of marine foraminifera, molluscs and brachiopods. As a proxy palaeobarometer of CO2, he uses stomatal index of fossil leaves of four closely related gymnosperms: Ginkgo, Lepidopteris, Tatarina and Rhachiphyllum. This palaeobotanical palaeobarometer is based on the observation that plant leaves have fewer stomates when atmospheric CO2 is high than when atmospheric CO2 is low. Taken together, these two proxies show strong support for CO2–temperature coupling over the past 300 Myr. Other proxy palaeobarometers—such as the carbon isotopic composition of palaeosols (Ekart et al . 1999; Tanner et al . 2001) and of marine phytoplankton compared with foraminifera (Pagani et al . 1999, 2000), which appear to indicate CO2–temperature uncoupling (Cowling 1999; Veizer et al . 2000)—may be compromised by isotopically unusual methane–clathrate dissociation events (Retallack 2001).
Temperature–CO2 coupling is demonstrated well by 300 Myr time-series for oxygen isotopic composition of marine biogenic carbonate and for CO2 concentrations derived from plant stomatal index, which show striking coincidences in peaks and valleys. The degree of correspondence is shown well by the cross-correlation of ages of palaeotemperature peaks inferred from oxygen isotopic values for marine biogenic carbonate and of CO2 peaks inferred from stomatal index (Retallack 2001). Thus times of warm palaeotemperature in the ocean from oxygen isotopic values correspond with times of high atmospheric CO2 from stomatal index.
Another important CO2 palaeobarometer has been derived from palaeosols, in which the carbon isotopic composition of pedogenic carbonate becomes heavier with higher partial pressure of atmospheric CO2 diffusing further into the soil to dilute isotopically light CO2 from respiration of plants and their debris. Unfortunately this palaeobarometer gives low CO2 levels (and some negative concentrations!) at times of known global warmth, such as the earliest Triassic (250 Ma), Early Jurassic (200– 190 Ma), Early Cretaceous (117 Ma) and Late Palaeocene (55 Ma (Ekart et al . 1999; Tanner et al . 2001)). Other isotopic studies demonstrate that each of these were times of massive dissociation of methane clathrates, flooding the atmosphere with CH4, another greenhouse gas. Methanotrophic CH4 is isotopically light (−110‰ and typically −60‰ δ13C), and this distinctive composition is an invaluable tracer of the temporal and geographic variation in abundance of CH4. Methane oxidizes to isotopically light CO2 within 7–24 yr, but observed perturbations of organic carbon isotopic composition attributed to methane hydrate dissociation lasted for hundreds of thousands of years. Carbon isotopic palaeobarometers of CO2 from palaeosols assume a gradient of isotopically light CO2 from respiration of soil organic matter to isotopically heavy atmospheric CO2. If atmospheric CO2 is isotopically light from methanogenic CH4 oxidation, then even with high atmospheric CO2, the isotopic composition of soil organic and inorganic carbon will be light and give the misleading impression that soil respiration was dominant and atmospheric CO2 low.
Other isotope-based paleobarometes such as that use the difference in carbon isotopic composition of particular organic compounds that are biomarkers for phytoplankton compared with the carbon isotopic composition of associated foraminifera in oligotrophic, deep marine sequences are likely to be compromised by episodes of methane hydrate dissociation and oxidation.
Important: an unexpected result of both the oxygen isotopic and stomatal index time-series is abrupt swings on geologically short time-scales. Even in the Early Tertiary, there were times of near-modern CO2 levels (Royer et al . 2001a), punctuated by periods of high CO2. It is thus a broad and unrealistic generalization to speak of an Early Tertiary greenhouse, let alone a Jurassic or Mesozoic greenhouse. Average CO2 levels were high during these time-spans, but CO2 maxima were sometimes very high and CO2 minima only a little higher than present CO2 levels.
Some of the CO2 highs revealed by stomatal index data are remarkably sharp, such as the Late Palaeocene thermal maximum (55 Ma), which is defined by a close temporal resolution of stomatal index measurements. The carbon isotopic excursion during the terminal Palaeocene is so profound that it can only be explained by dissociation of isotopically light CH4 from clathrate deposits. Methane clathrate dissociation events are also the only feasible interpretation of large carbon isotopic excursions during the earliest Triassic, Early Jurassic and Early Cretaceous. Some contribution of methane to the negative carbon isotopic anomaly is also likely at the Cretaceous–Tertiary boundary and during the Middle Miocene thermal optimum (herein), but in these cases the recorded magnitude of the negative carbon isotopic excursion is less than at the other times listed above, and can be explained by a variety of mechanisms, such as abrupt drops in productivity, as well as by methane release.
Other CO2 highs are less sharply focused and asymmetric in time, suggesting a short, sharp rise, followed by longer-term decay. These peaks may represent catastrophic disruption of the carbon cycle followed by long-term repair. Many of these were times of likely asteroid or comet impact, such as the terminal Permian, terminal Jurassic, and terminal Cretaceous. Many are times of flood basalt eruptions, such as the terminal Permian Siberian Traps, terminal Triassic Drakensburg Basalt, terminal Cretaceous Deccan Traps and Middle Miocene Columbia River Basalts. These dramatic forcings of the carbon cycle would be expected to leave transient residues in the atmosphere.
Some CO2 highs were times of geographical spread and adaptive radiation of tropical taxa to high latitudes, such as the Early Cretaceous adaptive radiation, and the rise to dominance of flowering plants and ornithischian dinosaurs. Middle Miocene (16 Ma) fossil floras of North America had unusually high percentages of exotic plants whose modern relatives live in Southeast Asia, which is the likely source of elephants, gelocid deer and other mammals immigrant to North America at the same time. Times of markedly elevated CO2 were also times of warm tropical marine palaeotemperature, as indicated by the oxygen isotopic composition of shellfish and foraminifera. Many CO2 highs follow times of elevated extinctions both on land and in the sea, such as the earliest Kazanian (mid-Permian), earliest Tatarian (Late Permian), earliest Griesbachian (earliest Triassic), Anisian (Middle Triassic), Carnian (Late Triassic), Hettangian (earliest Jurassic), Toarcian (Early Jurassic), Bathonian (Middle Jurassic), Berriasian (earliest Cretaceous), Aptian (Early Cretaceous), Cenomanian (mid-Cretaceous), Danian (earliest Palaeocene) and Ypresian (earliest Eocene). It thus appears that CO2 levels above 2000 ppmv and tropical marine palaeotemperatures above 25 ◦C are not compatible with high biological diversity. Or, as put more colourfully by Lovelock (2000), Earth's biota is sick when it runs a high temperature.
Also apparent from the data on stomatal index are extended periods during the Early Tertiary, mid-Cretaceous, Middle Jurassic and Early Triassic, when CO2 levels were moderately high (500–1000 ppmv). It is only in the Early Permian that values as low as modern are found. Perhaps some of these episodes can be explained by metamorphic or volcanic CO2 degassing rates higher than at present, but an alternative view is that they come from ecosystems with a consumer–producer ratio higher than at present. The carbon-oxidizing effects of termites and dinosaurs may explain generally higher levels of CO2 minima during the Mesozoic. This broader pattern is overridden by numerous catastrophic carbon-oxidizing events.
The numerous CO2 lows seen in stomatal index records reflect processes of carbon sequestration within biomass and by burial in carbonate and carbonaceous rocks. Two abiotic mechanisms proposed for carbon sequestration are mountain uplift and ocean-current upwelling. Uplift of the Himalaya beginning some 45 Ma has been argued to stimulate erosion and weathering in grasslands of the Tibetan plateau, which latter consumes CO2 by hydrolysis of newly exposed mineral grains and delivers more organic sediment and nutrients to the ocean. Thermal isolation of Antarctica by completion of the Circum-Antarctic current with opening of the Drake Passage at some 30 Ma increased upwelling of dissolved nutrients and phytoplankton productivity in the southwest Pacific Ocean, thus increasing carbon burial there. Neither of these events has quite the right timing to explain the principal CO2 drawdowns seen in the Tertiary stomatal index record during the Palaeocene (60 Ma), Late Eocene (34 Ma) and mid–late Miocene (15 Ma). Both mechanisms depend on biological amplification of soil carbon fixation and of oceanic carbon burial, which could have occurred for purely evolutionary reasons, for example by the evolution of lignin and thick peat in forests, and of sod and organic crumb peds in grassland soils. In Olsen's (1993) view, times of low CO2 were times when plants and plant-like microbes were thriving, reducing and fixing carbon by photosynthesis, at rates that exceeded carbon oxidation by actinobacteria, fungi or animals. The overall effect of life has been to cool our planet from long-term increase in solar radiation.
The global carbon cycle: storages (PgC) and fluxes (PgC/yr) estimated for the 1980s.
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The thick arrows denote the most important fluxes from the point of view of the contemporary CO2 balance of the atmosphere: gross primary production and respiration by the land biosphere, and physical air-sea exchange. These fluxes are approximately balanced each year, but imbalances can affect atmospheric CO2 concentration significantly over years to centuries. The thin arrows denote additional natural fluxes (dashed lines for fluxes of carbon as CaCO3), which are important on longer time-scales. The flux of 0.4 PgC/yr from atmospheric CO2 via plants to inert soil carbon is approximately balanced on a time-scale of several millenia by export of dissolved organic carbon (DOC) in rivers (Schlesinger, 1990). A further 0.4 PgC/yr flux of dissolved inorganic carbon (DIC) is derived from the weathering of CaCO3, which takes up CO2 from the atmosphere in a 1:1 ratio. These fluxes of DOC and DIC together comprise the river transport of 0.8 PgC/yr. In the ocean, the DOC from rivers is respired and released to the atmosphere, while CaCO3 production by marine organisms results in half of the DIC from rivers being returned to the atmosphere and half being buried in deep-sea sediments − which are the precursor of carbonate rocks. Also shown are processes with even longer time-scales: burial of organic matter as fossil organic carbon (including fossil fuels), and outgassing of CO2 through tectonic processes (volcanism). Emissions due to volcanism are estimated as 0.02 to 0.05 PgC/yr (Williams et al., 1992; Bickle, 1994).
Carbon cycling in the ocean
CO2 that dissolves in the ocean is found in three main forms (CO2, CO32-, HCO3-, the sum of which is DIC). DIC is transported in the ocean by physical and biological processes. Gross primary production (GPP) is the total amount of organic carbon produced by photosynthesis (estimate from Bender et al., 1994); net primary production (NPP) is what is what remains after autotrophic respiration, i.e., respiration by photosynthetic organisms (estimate from Falkowski et al., 1998). Sinking of DOC and particulate organic matter (POC) of biological origin results in a downward flux known as export production (estimate from Schlitzer, 2000). This organic matter is tranported and respired by non-photosynthetic organisms (heterotrophic respiration) and ultimately upwelled and returned to the atmosphere. Only a tiny fraction is buried in deep-sea sediments. Export of CaCO3 to the deep ocean is a smaller flux than total export production (0.4 PgC/yr) but about half of this carbon is buried as CaCO3 in sediments; the other half is dissolved at depth, and joins the pool of DIC (Milliman, 1993). Also shown are approximate fluxes for the shorter-term burial of organic carbon and CaCO3 in coastal sediments and the re-dissolution of a part of the buried CaCO3 from these sediments.
Carbon cycling on land
By contrast with the ocean, most carbon cycling through the land takes place locally within ecosystems. About half of GPP is respired by plants. The remainder (NPP) is approximately balanced by heterotrophic respiration with a smaller component of direct oxidation in fires (combustion). Through senescence of plant tissues, most of NPP joins the detritus pool; some detritus decomposes (i.e., is respired and returned to the atmosphere as CO2) quickly while some is converted to modified soil carbon, which decomposes more slowly. The small fraction of modified soil carbon that is further converted to compounds resistant to decomposition, and the small amount of black carbon produced in fires, constitute the "inert" carbon pool. It is likely that biological processes also consume much of the "inert" carbon as well but little is currently known about these processes. Estimates for soil carbon amounts are from Batjes (1996) and partitioning from Schimel et al. (1994) and Falloon et al. (1998). The estimate for the combustion flux is from Scholes and Andreae (2000). 'T' denotes the turnover time for different components of soil organic matter.
Human perturbation of natural carbon cycle
Fossil fuel burning and land-use change are the main anthropogenic processes that release CO2 to the atmosphere. Only a part of this CO2 stays in the atmosphere; the rest is taken up by the land (plants and soil) or by the ocean. These uptake components (depicted by red arrows) represent imbalances in the large natural two-way fluxes between atmosphere and ocean and between atmosphere and land.
An estimate of the human-induced carbon cycle in the 1990 (units are PgC yr-1). The carbon flows from fossil-fuel emissions to the atmosphere and the net carbon flows to the ocean and land are known with relatively high confidence. The partitioning of the net land sink between human activity and 'natural' carbon sinks is less certain, as is the partition between tropical and temperate regions.
References and further reading
- Ruddiman WF. Orbital insolation, ice volume, and greenhouse gases. Quaternary Science Reviews. Vol. 22, Issues 15–17, July–August 2003, pp 1597–1629
- Farquhar et al. The Carbon Cycle and Atmospheric Carbon Dioxide. (pdf available on-line)
- Malhi Y et al. Forests, carbon and global climate. Phil. Trans. R. Soc. Lond. A (2002) 360, 1567-1591. (.pdf available on-line)
- Why some gases are greenhouse gases, but most aren't, and some are stronger than others. (.pdf available on-line)
- Why some gases are greenhouse gases, but most aren't, and some are stronger than others. (.pdf available on-line)
- How do we know more CO2 is causing warming? Skeptical science
- Degens ET. The History of CO2. Environment International Vol. 2, pp. 401-408, 1979.
- Archer D et al. Atmospheric Lifetime of Fossil Fuel Carbon Dioxide. Annual Review of Earth and Planetary Sciences. Vol. 37: 117-134, May 2009.
- Solomon S et al. Irreversible climate change due to carbon dioxide emissions. Proc Natl Acad Sci U S A. 2009 Feb 10; 106(6): 1704–1709.
- Shaw GH. Earth's atmosphere – Hadean to early Proterozoic. Chemie der Erde 68 (2008) 235-264
- Retallack GJ. A 300-million-year record of atmospheric carbon dioxide from fossil plant cuticles. Nature. 2001 May 17;411(6835):287-90.
- Zhang YG, Pagani M, Liu Z, Bohaty SM, Deconto R. A 40-million-year history of atmospheric CO(2). Philos Trans A Math Phys Eng Sci. 2013 Sep 16;371(2001):20130096. (pdf available in the internet)
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