12  Terrestrial greenhouse gases

Author: Fabrice Lacroix

12.1 Atmospheric greenhouse gases and the Earth’s Energy Budget

Greenhouse gases are gases that absorb longwave radiation while letting shortwave radiation bypass, thus trapping heat within the atmosphere. The rise in concentration of such gases, such as CO₂, CH₄ and N₂O, has caused an energy surplus in the atmosphere and substantial warming of the climate since the preindustrial time period (Figure 12.1).

Understanding their sources and sinks, such the role of the terrestrial biosphere for greenhouse gas exchange, and how they are changing with climate and other anthropogenic dimensions of change, is key to understanding current and projecting future climate states. In this Chapter, we will look at overviews of greenhouse gas budgets, their evolution during interglacial cycles, and how they act to shape the climate.

Figure 12.1: Change in radiative forcing from 1750 to 2019 of greenhouse gases (CO₂, non-CO₂ greenhouse gases - broken down into contributions of N₂O, CH₄) versus other known forcing agents in the Earth system. Solid bars represent best estimates, and very likely (5–95%) ranges are given by error bars. From Forster et al. (2021)

12.2 Terrestrial sinks and sources of atmospheric greenhouse gases

In Chapter 3, it is explained in detail how increasing human-caused CO₂ emissions to the atmosphere, which are dominated by the direct burning of fossil fuel, are partly compensated by decadal terrestrial and oceanic sinks. According to our current knowledge, 50 /% of total emissions are absorbed by land and ocean, with the other half increasing atmospheric CO₂. Despite the role of the global terrestrial biosphere as a carbon sink, land also has emission pathways constituting as point sources to the atmosphere. These can arise from natural systems as well as managed lands. Land use change, in particular, is known to be an important contributor of CO₂ to the atmosphere since the preindustrial time period, owing to the clearing of natural forests to make way for agricultural land or urban landscapes (Friedlingstein et al. (2025), Chapter 10). Thawing of permafrost, disturbances such as fire, as well as drainage of wetlands can also cause substantial local sources of CO₂ emissions to the atmosphere. Overall however, unmanaged and covered land areas are still a stronger increasing sink of CO₂ globally, owing largely to the effect of CO₂ fertilization, which dominates over sources.

Land systems are also important for the budgets of the non-CO₂ greenhouse gases CH₄ and N₂O, which are also major drivers of radiative forcing changes since preindustrial (Figure 12.1). In contrary to the land exchange of CO₂, which constitutes of a net sink, land provide important source terms of CH₄ and N₂O to the atmosphere. Although atmospheric concentrations of CH₄ and N₂O are two-to-three orders of magnitude lower than concentrations of CO₂ (factor 100 to 1000), their contribution to anthropogenic climate change, measured by their Effective Radiative Forcing (ERF), is still significant. It is however not always straight forward to distinguish between unperturbed, natural fluxes from ecosystems, and emissions arising from the use and transformation of natural ecosystems for human activities.

12.3 Global CO₂ Budget

As illustrated in the Global Carbon Budget (Chapter 3), the terrestrial sphere plays an important role as a sink in the global CO₂ budget. Emissions due to land use change firstly account for a source term in the budget, accounting for around 10 % of total anthropogenic emissions, with the rest attributed to fossil fuel emissions. Secondly, land also takes up around 21% of total CO₂ emissions at the decadal timescale. This second term is dominant in terms of magnitude, meaning land is an overall sink with both effects included. The sink is of similar magnitude to its counterpart, the ocean sink (Chapter 13), but is much more variable from year to year. This is because land vegetation is more sentitive to yearly variability of the climate, compared to the carbonate chemistry-driven ocean sink. The terrestrial CO₂ sink is thought to be majorly driven by the CO₂-fertilization effect on the vegetation growth, which enhances carbon uptake in model simulations (Friedlingstein et al. (2025)). This effect of CO₂ fertilization may explain a large part of observed greening trends around the globe, and has been demonstrated in elevated CO₂ experiments (Walker et al. (2015)), as well as tree ring records (Hubau et al. (2020)). Climate change-driven effects can in turn have different signs of change on the terrestrial sink. It largely acts towards reducing the land carbon sink, especially through the enhancement of soil emissions and decreased plant growth in the tropics (Friedlingstein et al. (2025)). However, in high latitudes, where low temperatures are an important limiting factor for vegetation growth, terrestrial CO₄ uptake are increasing due to warmer conditions, as well as due to longer growing seasons (Berner et al. (2020)). More details on the global carbon budget and spatial CO₂ patterns can be found in its dedicated Chapter Chapter 3.

Figure 12.2: Decomposition of spatial global CO₂ uptake patterns of the land biosphere into contributions from atmospheric CO₂ concentrations alone, thus of the CO₂ fertilization effect (panels c and d), and climate-driven changes alone (panel a and b). From Friedlingstein et al. (2025) .

12.4 Global Methane (CH₄) Budget

Generally, methane sources to the atmosphere occur due to fossil fuel burning, production in wet soils and freshwaters, as well as through agricultural activities. Methane emissions arise both from natural processes (dark cyan in Figure 12.3), contributing to its atmospheric concentrations, but its concentrations in the atmosphere have also substantially increased due direct human activities (anthropogenic perturbation), such as fossil fuel, emission from landfills, and livestock. Climate warming also is thought to enhance methane emissions from natural ecosystems (warming-induced emissions), although these changes are very hard to measure and quantify. In contrast to CO₂, which essentially does not react in the atmosphere, methane is broken down in the atmosphere, primarily through a tropospheric reaction with hydroxyl, as well as through further other minor tropospheric and stratospheric consumption processes (dark red in Figure 12.3). Atmospheric loss is thus the dominant component of methane removal in the Earth system. A certain amount of methane is also sequestered on land through its oxidation in soils. This is usually methane that was produced in water-satured part of the soil.

Natural land systems emits methane in wetlands, freshwater systems, and geological processes. For the former two, these sources of methane, which arise in conditions where oxygen levels are very low (anoxic conditions), are usually the residual of production of methane within the ecosystem and its oxidation within the soil or water column. As depicted in Figure 12.3, wetlands and wet soils dominate emissions with 159 Tg CH₄ yr^-1 emitted per year, mainly through production by methanogen bacteria. Spatially, natural land-induced methane emissions are particularly high in tropical wetlands, where wet and warm conditions provide ideal conditions for microbial production. Globally-significant emissions also occur in carbon-rich wetlands (peatlands) and permafrost-affected soils of higher latitudes. Further global natural land emissions of methane arise in freshwater systems, particularly lakes with sediments rich in organic matter, which emit an average of 112 Tg CH₄ yr^-1 to the surface through the water column. Here again, low latitude regions form dominant areas for freshwater emissions due to warm conditions, while high latitude freshwaters can also account for globally-significant sources due to large carbon concentrations in rivers, streams and lakes. Terrestrial geological processes furthermore emit 45 Tg CH₄ annually from the Earth’s crust through faults and fractured rocks and seeps. Termites also contribute to methane emissions due to the byproduction during their digestion of plant material. Besides land sources, the ocean also contributes a smaller share natural, with only 12 Tg CH₄ yr^-1 on average, through biogenic, geological and hydrate emissions from coastal and open ocean.

Changes in soil conditions (warming, soil moisture change), in freshwater temperatures, as well as warming of ocean waters and changes in circulation can affect the production of methane and its oxygenation, although these changes remain strongly uncertain at the global scale. Particularly, thawing of permafrost and changes in wetland areas are likely increasing CH₄ emissions. Thawing of the permafrost, for instance, exposes large amounts of organic carbon to microbial degradation in low oxygen conditions. Previously, this carbon had been immobilized for thousands of years at freezing temperatures. The potential release of CH₄ following the thawing process may cause a positive (warming) climate feedback with future warming, which could have significant implications for the global climate (see Chapter 9).

Figure 12.3: Global methane (CH₄) budget (2010–2020). Figure basis taken from Canadell et al. (2021) and updated with fluxes from Saunois et al. (2025).

NoteSoil CH₄ Emissions

CH₄ emissions are produced by microbes (“methanogens”) under wet and anaerobic (low oxygen) conditions, resulting in a soil sink of CH₄. This production process occurs in completely saturated soil parts, or in completely water-saturated wetlands. Methane production can take place through two distinct processes:

CO₂ reduction

CO₂ + 4H₂ -> CH₄ + H₂O

Acetate fermentation

CH₃COOH + 4H₂ -> CH₄ + H₂O

The emissions can then reach the surface through diffusion or ebullition (bubble formation and ascent).

The produced CH₄ can be oxidized within the soil (often in the shallow part of the soil), following the equation:

Methane oxidation

CH₄ + O₂ -> CO₂ + 2H₂

Figure 12.4: Schematic figure showing the methane formation pathway in water-saturated soils (panel a) and in non-saturated soils (panel b). From Riley et al. (2011).

Direct anthropogenic sources of methane primarily arise from fossil fuel extraction and processing, livestock, waste management, and rice cultivation. Fossil fuel extraction is estimated to contribute 120 Tg CH₄ yr^-1, with significant emissions arising from countries such as China, the United States, and Russia, which are the leading producers of coal, oil, and gas, globally. Livestock, particularly ruminants, are the largest source linked to agricultural activities, emitting 112 Tg CH₄ yr^-1 on average through enteric fermentation and manure management. Similarly, anaerobic decomposition occurs in landfills and during organic waste treatment, as well as in flooded rice paddies. Additionally, methane is released during incomplete combustion of organic material in biomass burning. Biomass burning contributes to methane emissions during the burning of crop residues, savannah fires, forest fires, and stoves.

12.5 Global nitrous oxide (N₂O) budget and terrestrial sources

Current N₂O sources to the atmosphere consist of sources from land and oceanic natural systems, as well as a number of anthropogenic sources such as agricultural processes. Due to these anthropogenic sources, atmospheric concentrations of N₂O have also experienced a substantial rise since the start of the industrial time period. As for methane, the loss of N₂O in the Earth system occurs to a large part by atmospheric degradation. In contrast to CH₄, this breakdown happens in the stratosphere instead of troposphere.

Global natural sources of N₂O include soils covered by unmanaged vegetation, oxygen-deprived zones of the open ocean, as well as smaller contributions from inland water bodies and atmospheric chemistry (Figure 12.5). Land ecosystems produce N₂O through microbial nitrification and denitrification processes in the soil. Nitrification, which uses NO₃^- as a substrate and energy source, occurs under oxic conditions, whereas N₂O emissions occur under incomplete denitrification in anoxic conditions. Denitrification, which is a series of processes, bacteria use NO₃^- as an oxidant to break down organic matter for energy. Natural sources of land derived N₂O thereby add up to 4.9-6.5 TgN yr^-1 through microbial nitrification and denitrification, as shown in Figure 12.6.

Other than terrestrial sources, the open ocean emits 2.5-4.3 TgN yr^-1 of N₂O to the atmosphere, largely through nitrification in productive but oxygen-rich waters, and denitrification in subsurface low-oxygen waters, as well as in coastal sediments.

Figure 12.5: Global nitrous oxide (N₂O) budget (2007–2016). The atmospheric stock is calculated from mean N₂O concentration, multiplying a factor of 4.79 ± 0.05 Tg ppb^–1. Pool sizes for the other reservoirs are largely unknown. Figure and caption taken from Canadell et al. (2021).

Of the anthropogenic N₂O emissions, fossil fuel burning is only here a minor emitor of N₂O. Agriculture is the most important source of emissions to the atmosphere, contributing 2.5–5.8 TgN yr^-1 through the extensive use of synthetic fertilizers and manure on croplands and pastures, as well as from manure management and aquaculture. Fertilizers and manure enrich the soil nutrition for plants but also increase N₂O emissions from the excess strongly reactive soil inorganic nitrogen that is not absorbed by plants. A further source of N₂O arises through the enhancement of reactive nitrogen in soils due to nitrogen deposition. Non-agricultural sectors also contribute to anthropogenic N₂O emissions, particularly from industrial activities such as fossil fuel combustion and chemical processing, and wastewater treatment. Industrial N₂O emissions, primarily from nitric and adipic acid production, have decreased in North America and Europe since the 1990s due to the adoption of abatement technologies. Additionally, inland water bodies, estuaries, and nitrogen deposition from atmospheric pollution contribute relatively small amounts of N₂O. Biomass burning, including natural and anthropogenic fires, also contributes minor amounts of N₂O to the atmosphere.

NoteSoil N₂O Emissions

During the soil N cycle (see Chapter 11), N₂O emissions can be produced during two distinct processes. Firstly, they can be produced during nitrification:

2NH₃ + 2O₂ -> NO₃^- + H⁺

which produces N₂O as a byproduct.

Secondly, it is also an intermediate product in denitrification, from the transformation of nitrate (NO₃), which is used as an oxidant, to nitrogen gas (N₂). This process takes place under wet and anaerobic (low oxygen) conditions:

NO₃^- -> NO₂^- -> NO -> N₂O

This N₂O can however also further react in the complete denitrification process, which reduces the amount of N₂O produced in denitrification:

2N₂O -> N₂ + O₂

Figure 12.6: Schematic illustration of the soil nitrogen cycle and nitrous oxide (N₂O) emissions during nitrification and denitrification processes.

12.6 History of atmospheric concentrations of CH₄ and N₂O

CO₂, CH₄ and N₂O concentrations have varied strongly over glacial-interglacial cycles, which are reconstructed from ice core concentrations. However, they all have rapidly risen in the anthropocene, especially sharply since industrialisation, with both CH₄ and N₂O reaching concentrations far beyond levels ever recorded during the last 800,000 years.

During glacial periods, atmospheric CO₂ and CH₄ concentrations were substantially lower than during warmer (interglacial) periods, when concentrations (Figure 12.7). Lower concentrations in glacial period owes to multiple reasons. While ice age cold periods are initially driven by periodic orbital settings (eccentricity, precession, tilt), that superimpose to cool the global temperate, it also sets off re-enforcing feedbacks of reduced atmospheric concentrations, which then lead to further cooling. Firstly, a colder ocean stores more carbon as CO₂ than a warmer ocean, which reduces atmospheric CO₂. Colder temperatures and reduced wetland coverage (due to freezing) further lead to the reduction of CO₂, CH₄ and N₂O emissions from soils. Lastly, the increased formation of glaciers and ice sheets reduces the planetary albedo, which further cools the climate. This feedback loops coolsa the planet until orbital settings become favorable to warming.

The transitions between glacial and interglacial periods in contrast were typically characterised by a sharp increase in greenhouse gas concentrations, reflecting CO₂ emissions from a warming ocean, as well as increases in natural CH₄ and N₂O sources, such as from soil warming, increasing wetlands and permafrost thawing. For instance, N₂O concentrations varied between about 195 ppb during glacial periods and up to 290 ppb during interglacial periods. These variations in atmospheric N₂O concentrations were related to varying sources from soils and marine emission changes in response to changes in ocean currents and oxygen availability (Joos and Spahni 2008). Also here, a reduction of ice sheet coverage reduced the planetary albedo, which further warmed the climate, leading to a warm period with high atmospheric greenhouse gas concentrations.

Figure 12.7: Figure 12.4: Atmospheric concentrations of CO₂, CH₄ and N₂O in air bubbles and clathrate crystals in ice cores (800,000 BCE to 1990 CE). Note the variable x-axis range and tick mark intervals for the three columns. Ice core data is over-plotted by atmospheric observations from 1958 to present for CO₂, from 1984 for CH₄ and from 1994 for N₂O. The time-integrated, millennial-scale linear growth rates for different time periods (800,000–0 BCE, 0–1900 CE and 1900–2017 CE) are given in each panel. For the BCE period, mean rise and fall rates are calculated for the individual slopes between the peaks (interglacials) and troughs (glacial periods), which are given in the panels in left column. Figure from Canadell et al. (2021).

The last glacial maximum occurred around 20,000 years ago. Since then, both CH₄ and N₂O concentrations have increased first to typical interglacial levels during the pre-industrial Holocene (between around 11 kyr BP to 1700 CE), and then in a rapidly accelerated fashion to current levels of around 1800 ppb CH₄ and around 330 ppb N₂O. The growth rates during the industrial period are about 500 times faster than ever recorded during glacial-interglacial cycles. This underlines the significant contribution of anthropogenic activities since the Industrial Revolution.

12.7 Atmospheric sinks of CH₄ and N₂O

As illustrated in Figure 12.3, methane is primarily removed from the atmosphere through chemical loss in the troposphere. This process, which consists of oxidation of methane by hydroxyl radicals (OH^-), removes 536 Tg CH₄ yr^-1 annually:

\[ CH_{4} + OH^{. } = CH_{3} + H_{2}O \]

Other minor sinks for methane include soil uptake through bacterial oxidation, loss in the stratosphere via reactions with chlorine (Cl) and excited oxygen atoms (O(¹D)), and loss through tropospheric chlorine reactions. Since the reaction with OH consists of around 90 % of CH₄ removal, the concentration of OH, particularly in the troposphere, is a dominant determinant of the rate of CH₄ removal.

By contrast, as illustrated in Figure 12.5, nitrous oxide is removed in the stratosphere. Photolysis, whereby N₂O molecules are broken down by ultraviolet (UV) radiation, is by far the major process of N₂O removal.

\[ N_{2}O + hv = N_{2} + O \]

The resulting oxygen atoms, O(¹D) radicals, also oxidize N₂O and break it down. Stratospheric environmental factors, such as temperature and UV radiation levels, play a major role in N₂O destruction. Warming could potentially accelerate stratospheric circulation, affecting the distribution and degradation of N₂O. Furthermore, changes in stratospheric ozone concentrations can influence the rate at which N₂O is broken down.

Sink rates are proportional to the respective atmospheric concentration for both CH₄ and N₂O is that, at the first order. The atmospheric breakdown of CH₄ and N₂O, scales proportionally with number of CH4 and N₂O molecules exposed to breakdown processes.

12.8 Atmospheric lifetime of CH₄ and N₂O

Atmospheric processes are key sinks of both CH₄ and N₂O in the Earth system, largely compensating for their present-day sources. These atmospheric sink processes for CH₄ and N₂O (Section 12.7), which are approximately proportional to their concentration, and their emission sources, determine the atmospheric mass balance of CH₄ and N₂O: \[ \frac{\mathrm{d}M(t)}{\mathrm{d}t} = E(t) - k * M(t) \;. \tag{12.1}\]

\(M(t)\) is the atmospheric mass of either CH₄ or N₂O at time \(t\), and \(E(t)\) are the respective emissions. k is the degradation rate, and an atmospheric lifetime, the turnover time (in years), of the greenhouse gas can be defined from the equation:

\[ \tau = \frac{1}{k} \tag{12.2}\]

Methane (CH₄) is a relatively short-lived GHG with a \(\tau\) of about 11.8 ± 1.8 years (Table 12.1), while nitrous oxide (N₂O) is considered a long-lived GHG, persisting in the atmosphere for around 109 ± 10 years (Forster et al. 2021). This means that methane has a shorter atmospheric residence time compared to the longer-residing N₂O, allowing it to reach a steady state more quickly due to its higher rate of removal. In contrast, N₂O may accumulate in the atmosphere over longer time scales, remaining for more than a century before it is removed. CO₂ on the other hand, can have strongly differing atmospheric lifetimes. For one, emissions of CO₂ are rapidly taken up by land or ocean within a few years. The remaining CO₂ in the atmosphere however has lifetime of centuries due to the lack of removal processes in the atmosphere. The remaining is only very slowly taken up by land and ocean.

We can further use the combined equations to derive a mass change based on \(\tau\):

\[ \frac{\mathrm{d}M(t)}{\mathrm{d}t} = E(t) - \frac{1}{\tau} * M(t) \tag{12.3}\]

See Section 3.2 for a more detailed introduction of the concept of the turnover time and the 1st-order decay model.

Finally, setting \[ \frac{\mathrm{d}M(t)}{\mathrm{d}t} = 0 \]:

we can derive under fixed emissions E:

\[ M_{steadystate} = \frac{E}{k} \tag{12.4}\]

NoteGlobal warming potential

Differences in atmospheric residence times of difference greenhouse gases have implications for their warming effects over time. Greenhouse gas warming effects are often measured relative to effects of carbon dioxide (CO₂), thus in terms of the same mass (kilograms of CO₂ equivalent). This results in a dimensionless factor known as the Global Warming Potential (GWP), which is always relative to CO₂. The GWP metric quantifies the cumulative warming effect (radiative efficiency) of the emission of 1 kilogram of a greenhouse gas compared to the emission of 1 kilogram of CO₂ and considers its removal from the atmosphere over a specified time horizon, commonly 20 years (GWP20), 100 years (GWP100), or 500 years (GWP500). Thus, the GWP considers both the radiative efficiency of the GHG and its atmospheric lifetime.

It is important to note also that the GWP of the short-lived GHG CH₄ is much larger at the 20-year time scale, than at longer time scales. This owes to the rapid decay of atmospheric CH₄, while part of the CO₂ remains airborne for longer (a considerable fraction remains in the atmosphere for centuries). In contrast to CH₄, N₂O has a similar GWP at the 20 and 100 year time scale because its atmospheric concentration decline unfolds at a similar rate than that of CO₂ over the first 100 years.

Table 12.1: Emissions metrics for selected species: global warming potential (GWP). Table from Forster et al. (2021).

Species	Lifetime (Years)	Radiative Efficiency (W m⁻² ppb⁻¹)	GWP-20	GWP-100	GWP-500
CO2	Multiple	1.33 ± 0.16 ×10⁻⁵	1.0	1.0	1.0
CH4	11.8 ± 1.8	5.7 ± 1.4 ×10⁻⁴	79.7 ± 25.8	27.0 ± 11	7.2 ± 3.8
N2O	109 ± 10	2.8 ± 1.1 ×10⁻³	273 ± 118	273 ± 130	130 ± 64

The emission-concentration relationships for CO₂, CH₄, and N₂O differ significantly due to their different sources, atmospheric lifetimes, and the mechanisms by which atmospheric concentrations are reduced. CO₂ has multiple lifetimes as it cycles between different pools. Once emitted, atmospheric CO₂ is not actually broken down in the atmosphere (no chemical destruction), but is gradually redistributed among the different spheres in the Earth system - a process that takes hundreds to thousands of years. Critically, a portion of CO₂ remains in the atmosphere at millennial time scales, leading to a continuous buildup in atmospheric concentration over time unless emissions are reduced to zero. In contrast, CH₄ has a shorter atmospheric lifetime of decades and is removed by chemical reactions in atmosphere. After a cessation of emissions, atmospheric CH₄ concentrations decline exponentially towards zero. Analogously, when CH₄ emissions are stabilised, concentrations are stabilised as well, while CO₂ concentrations continue rising under positive non-zero emissions. The same dynamics as for CH₄ apply to N₂O, albeit considering its longer atmospheric lifetime, N₂O has much slower dynamics and response time scales of concentration changes in response to emission changes.

12.9 Future evolution of CO2, CH₄ and N₂O in future scenarios and the transient climate response to emissions

Rising atmospheric CO₂ concentrations, climate change, increasing reactive nitrogen inputs (atmospheric deposition and N fertiliser applications on croplands) all will continue to contribute to changes in terrestrial emissions CH₄ and N₂O in future scenarios, according to the newest models. The CO₂ fertilization effect on vegetation growth will likely persevere with further rising atmospheric CO2. As a consequence of warming, some models suggest that CO₂, N₂O and CH₄ production from natural systems will increase in a high-emission scenario, but can be stabilized in scenarios with lower warmng (Stocker et al. 2013). Additionally, future inputs reactive N inputs from fertilizers application and atmospheric deposition could increase substrate availability for denitrification, leading to higher N₂O emissions. However, climate change also alters soil moisture and oxygen availability, which is a control on both N₂O and CH₄ emissions from soils. Plant productivity enhancements under rising CO₂ and climate change tend to deplete inorganic soil N pools and may thus reduce denitrification rates and N₂O production. These net effects arising from multiple changing drivers are simulated differently by different models, thus leading to considerable uncertainty in future scenario projections (Zaehle et al. 2011; Stocker et al. 2013).

A concept that has arose to very simply estimate future warming based on future emission scenarios is that of the transient climate response to emissions Figure 12.8. In this framework, the warming observed and simulated by models (Y-axis) is compared to cumulative CO₂ emissions on the X-axis. The figure shows an almost linear evolution of warming to cumulative CO₂ emissions, with small deviations explained by non-CO₂ greenhouse gas emissions, as well as changes in other forcing agents (aerosols, CFCs). While a saturation of the greenhouse gas sink in both ocean and land is likely increasingly in high emission scenarios - meaning that the ocean and land take up a declining fraction of emissions human produce - compensating feedbacks from the atmospheric heat budget is thought to largely compensate the decreasing efficiency in atmospheric CO₂ removal. This is however still an area of ongoing research.

Figure 12.8: Transient climate response to emissions. Figure from Canadell et al. (2021).

CautionExercise

1. Projected fossil fuel emissions for the year 2025 were estimated to be 11.8 GtC yr_-1, marking a new record value. According to fractions of uptake in @ref(sec-globalcarbonbudget), how much of these emissions remain in the atmosphere, and how much are taken up by land and ocean on a decadal timescale?
2. Assume constant CH₄ and N₂O emissions of 610 MtCH4 yr^-1 and 17 Mt N₂O yr^-1. What is the steady state level of the gases in the atmosphere?
3. The best estimate of the transient climate response to emissions is currently of around 1.5 °C / 1000 Gigatons CO₂. Assuming scenarios of future emissions of 270 (SSP1-26), 750 (SSP2-45), 1438(SSP3-70)and 2103 (SSP5-85) of Gigatons C of CO₂, warming would you expect in year 2100?
4. Explain why the Global Warming Potential of greenhouse gases differ and change over time?
5. Considering we need to severely reduce greenhouse gas emission to meet climate goals in year 2050. Which greenhouse gas would be most efficient to reduce, from an Earth system perspective only, and why?

Berner, Logan T., Richard Massey, Patrick Jantz, Bruce C. Forbes, Marc Macias-Fauria, Isla Myers-Smith, Timo Kumpula, et al. 2020. “Summer Warming Explains Widespread but Not Uniform Greening in the Arctic Tundra Biome.” Nature Communications 11 (1): 4621. https://doi.org/10.1038/s41467-020-18479-5.

Canadell, J. G., P. M. S. Monteiro, M. H. Costa, L. Cotrim da Cunha, P. M. Cox, A. V. Eliseev, S. Henson, et al. 2021. “Global Carbon and Other Biogeochemical Cycles and Feedbacks.” Book Section. In Climate Change 2021: The Physical Science Basis. Contribution of Working Group i to the Sixth Assessment Report of the Intergovernmental Panel on Climate Change, edited by V. Masson-Delmotte, P. Zhai, A. Pirani, S. L. Connors, C. Péan, S. Berger, N. Caud, et al., 673–816. Cambridge, United Kingdom; New York, NY, USA: Cambridge University Press. https://doi.org/10.1017/9781009157896.007.

Forster, P., T. Storelvmo, K. Armour, W. Collins, J.-L. Dufresne, D. Frame, D. J. Lunt, et al. 2021. “The Earth’s Energy Budget, Climate Feedbacks, and Climate Sensitivity.” Book Section. In Climate Change 2021: The Physical Science Basis. Contribution of Working Group i to the Sixth Assessment Report of the Intergovernmental Panel on Climate Change, edited by V. Masson-Delmotte, P. Zhai, A. Pirani, S. L. Connors, C. Péan, S. Berger, N. Caud, et al., 923–1054. Cambridge, UK; New York, NY, USA: Cambridge University Press. https://doi.org/10.1017/9781009157896.009.

Friedlingstein, Pierre, Michael O’Sullivan, Matthew W. Jones, Robbie M. Andrew, Judith Hauck, Peter Landschützer, Corinne Le Quéré, et al. 2025. “Global Carbon Budget 2024.” Earth System Science Data 17 (3): 965–1039. https://doi.org/10.5194/essd-17-965-2025.

Hubau, Wannes, Simon L Lewis, Oliver L Phillips, Kofi Affum-Baffoe, Hans Beeckman, Aida Cunı́-Sanchez, Armandu K Daniels, et al. 2020. “Asynchronous Carbon Sink Saturation in African and Amazonian Tropical Forests.” Nature 579 (7797): 80–87.

Joos, Fortunat, and Renato Spahni. 2008. “Rates of Change in Natural and Anthropogenic Radiative Forcing over the Past 20,000 Years.” Proceedings of the National Academy of Sciences 105 (5): 1425–30. https://doi.org/10.1073/pnas.0707386105.

Riley, W. J., Z. M. Subin, D. M. Lawrence, S. C. Swenson, M. S. Torn, L. Meng, N. M. Mahowald, and P. Hess. 2011. “Barriers to Predicting Changes in Global Terrestrial Methane Fluxes: Analyses Using CLM4Me, a Methane Biogeochemistry Model Integrated in CESM.” Biogeosciences 8 (7): 1925–53. https://doi.org/10.5194/bg-8-1925-2011.

Saunois, M., A. Martinez, B. Poulter, Z. Zhang, P. A. Raymond, P. Regnier, J. G. Canadell, et al. 2025. “Global Methane Budget 2000–2020.” Earth System Science Data 17 (5): 1873–1958. https://doi.org/10.5194/essd-17-1873-2025.

Stocker, Benjamin, Raphael Roth, Fortunat Joos, Renato Spahni, Marco Steinacher, Soenke Zaehle, Lex Bouwman, Xu-Ri, and Iain Colin Prentice. 2013. “Multiple Greenhouse-Gas Feedbacks from the Land Biosphere Under Future Climate Change Scenarios.” Nature Climate Change 3 (7): 666–72. https://doi.org/10.1038/nclimate1864.

Walker, Anthony P, Sönke Zaehle, Belinda E Medlyn, Martin G De Kauwe, Shinichi Asao, Thomas Hickler, William Parton, et al. 2015. “Predicting Long-Term Carbon Sequestration in Response to CO2 Enrichment: How and Why Do Current Ecosystem Models Differ?” Global Biogeochemical Cycles 29 (4): 476–95.

Zaehle, Sönke, Philippe Ciais, Andrew D. Friend, and Vincent Prieur. 2011. “Carbon Benefits of Anthropogenic Reactive Nitrogen Offset by Nitrous Oxide Emissions.” Nature Geoscience 4 (9): 601–5. https://doi.org/10.1038/ngeo1207.