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, since the preindustrial time period has caused an energy surplus in the atmosphere and substantial warming of the climate (Figure 12.1). Understanding their sources and sinks, for which terrestrial sources and sinks play an important role, and how these are changing with climate and other anthropogenic change, is key to understand current and project future warming of the planet due to human activities.

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

Chapter 3 that increasing anthropogenic CO₂ to the atmosphere, which are dominated by the direct burning of fossil fuel, are partly compensated by terrestrial and oceanic sinks. Despite this sink, land also has emissionpathways from natural systems as well as managed lands, which can be sources of CO₂, CH₄ and N₂O to the atmosphere. Land use change, in particular, is known to be an important contributor of CO₂ to the atmosphere since the preindustrial time period (Friedlingstein et al. (2025)). Thawing of permafrost, disturbances such as fire, as well as drainage of wetlands can cause substantial local sources of CO₂ emissions to the atmosphere. Overall however, unmanaged and covered land areas are still an increasing sink globally of CO₂ owing largely to the effect of CO₂ fertilization, which dominates over sources.

Terrestrial systems are important for the budgets of the non-CO₂ greenhouse gases CH₄ and N₂O, which are amongst the most potent radiative forcing agents (Figure 12.1), by providing important source terms to the atmosphere. Although atmospheric concentrations of CH₄ and N₂O are two to three orders of magnitude lower than concentrations of CO₂ (factor 100-1000), their contribution to anthropogenic climate change, measured by their effective radiative forcing, in particular, is still significant due to their high global warming potential (see Section 12.8 below). Amongst them, methane (CH₄) has the highest effective radiative forcing (ERF) since 1700 with 0.54 [0.43-0.65] W m^-2, which is comparable in magnitude to the ERF of halogens and ozone. Nitrous oxide (N₂O) also has a sizable contribution to anthropogenic climate change with an ERF of 0.21 [0.18-0.24] W m^-2, amounting to roughly 10% of the ERF of CO₂, which is 2.16 [1.90-2.41] W m^-2. Both CH₄ and N₂O are emitted from a large number of natural and anthropogenic sources. It is however currently difficult 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 shown in the Global Carbon Budget (Chapter 3), the terrestrial sphere plays an important role in the global CO₂ budget: The terrestrial Earth system component accounts for part of global emissions from land use change, but it has taken up around 21% of total emissions, absorbing atmospheric CO₂. This sink is of similar magnitude to its counterpart: the ocean sink (Chapter 13). The terrestrial CO₂ sink is thought to be majorly driven by the CO₂-fertilization effect on the vegetation growth and carbon uptake based on 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 have in turn different signs of change on the terrestrial sink. It largely acts towards reducing this sink, especially through the enhancement of soil emissions and decreased plant growth in the tropics (Friedlingstein et al. (2025)). However, in the high latitudes, warming is increasing terrestrial CO2 uptake by providing better (warmer) plant growth conditions, as well as longer growing seasons ((berner?)) More details on the global carbon budget and spatial CO₂ patterns can be found in its dedicated Chapter Chapter 3.

12.4 Global Methane (CH₄) Budget

Generally, methane is produced through fossil fuel burning, in wet soils and freshwaters, as well as through agricultural activities. Methane emissions arise from both natural processes (dark cyan in Figure 12.2), contributing to its atmospheric concentrations, but concentrations have been increasing due to enhanced emissions from human activities, as well as through human-caused warming. The atmosphere is a sink of methane primarily through tropospheric hydroxyl loss, and other minor tropospheric and stratospheric consumption processes (dark red in Figure 12.2). A certain amount of methane is also sequestered through its oxidation in soils.

The largest natural terrestrial sources of methane include emissions from wetlands, freshwater systems, and geological processes. For the former two, these sources of methane, which arise in conditions where oxygen levels are very low (anaerobic conditions), are thereby the residual of production of methane within the ecosystem and its oxidation within the soil or water column. As depicted in Figure 12.2, wetlands contribute 102-182 Tg CH₄ yr^-1, mainly through microbial production under anaerobic conditions. Methane emissions are particularly high in tropical wetlands, where red and warm conditions provide ideal production conditions. There are also substantial emissions arising in carbon-rich wetlands of higher latitudes. Further global emissions of methane arise from freshwater systems, particularly lakes with sediments rich in organic matter, which emit 117-212 Tg CH₄ yr^-1 via ebullition (bubble formation) and diffusion to the surface through the water column. Low latitude areas are also dominant for freshwater emissions due to warm conditions, while high latitude freshwaters can also account for globally-significant sources due to large carbon concentrations in the water. Terrestrial geological processes furthermore emit 18–65 Tg CH₄ annually from the Earth’s crust through faults and fractured rocks and seeps. The ocean contributes a smaller share, with only 4–10 Tg CH₄ yr^-1, through biogenic, geological and hydrate emissions from coastal and open ocean, while termites, an insect infraorder, release around 9 Tg CH₄ yr^-1.

Changes in soil conditions (warming, soil moisture change), in freshwater temperatures, as well as warming of ocean waters and changes in circulation can affect methane production and oxygation, although these changes remain strongly uncertain at the global scale. Particularly, the thawing of permafrost and changes in wetland areas may be causing changes in CH₄ emissions. Thawing of the permafrost, for instances, is exposing large amounts of organic carbon to degratation. Previously, this carbon had been immobilized for thausends of years at freezing temperatures. The potential release of CH₄ following the thawing process may cause a positive (warming) climate feedback with future warming, that could have significant impacts on the global climate (see Chapter 9).

Figure 12.2: Global methane (CH₄) budget (2008–2017). The atmospheric stock is calculated from mean CH₄ concentration, multiplying a factor of 2.75 ± 0.015 Tg ppb^–1, which accounts for the uncertainties in global mean CH₄. Figure and caption taken from Canadell et al. (2021).

NoteSoil CH₄ Emissions

CH₄ emissions are produced by microbes (“methanogens”) under wet and anaerobic (low oxygen) conditions. This process takes place in completely saturated soils, or wetlands. Production follows the net equation:

CO₂ + H₂ -> CH₄ + H₂O

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

CH₄ + O₂ -> CO₂ + 2H₂

Figure 12.3: 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 114–116 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 106–115 Tg CH₄ yr^-1 through enteric fermentation and manure management in anaerobic conditions. Similarly, anaerobic decomposition occurs in landfills and during organic waste treatment, as well as in flooded rice paddies (25-37 Tg CH₄ yr^-1). Additionally, methane is released during incomplete combustion of organic material in biomass burning. Biomass burning contributes 22-39 Tg CH₄ yr^-1, with emissions coming from burning crop residues, savannah fires, forest fires, and stoves. However, distinguishing between natural and human-caused fires is complicated by a variety of factors, including weather conditions, vegetation types and land management practices.

12.5 Global nitrous oxide (N₂O) budget and terrestrial sources

Current N₂O sources to the atmosphere consist of natural terrestrial and oceanic sources, as well as a number of anthropogenic sources occuring from the burning of fossil fuel, production in sewage water, as well as in 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 ((fig?)). The loss of N2O in the Earth system occurs through stratospheric oxydation.

Global natural sources of nitrous oxide (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 12.4). Terrestrial ecosystems produce N₂O through microbial nitrification and incomplete denitrification, which both are redox processes occuring in soils. Nitrification, which uses NO₃^- as a substrate, occurs under oxic conditions, whereas N₂O emissions occur under incomplete denitrification, which uses NO₃^- as an oxidant, under largely anoxic conditions. Natural sources of N₂O thereby add up to 4.9-6.5 TgN yr^-1 through microbial nitrification and denitrification, as shown in Figure 12.5.

Other than terrestrial sources, fossil fuel burning produces a source of 2.5-4.3 TgN yr^-1 to the atmosphere, while the open ocean emits 2.5-4.3 TgN yr^-1, largely through nitrification in productive but oxygen-rich waters, and denitrification in subsurfacce oxygen minimum zones, as well as in coastal sediments.

Figure 12.4: 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, agriculture is the most important source, contributing 2.5–5.8 TgN yr^-1 due to 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 soil inorganic nitrogen that is not absorbed by plants. A further -0.6-1.1 TgN yr^-1 could be added to soils in natural ecosystems influenced by human perturbation of nitrogen deposition. Non-agricultural sectors also contribute to anthropogenic N₂O emissions, particularly from industrial activities such as fossil fuel combustion and chemical processing (0.8-1.1 TgN yr^-1), and wastewater treatment (0.2-0.5 TgN yr^-1). 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 0.5-0.8 TgN yr^-1.

NoteSoil N₂O Emissions

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

2NH₃ + 2O₂ -> NO₃^- + H⁺

Secondly, it is also a byproduct of denitrification, where nitrate (NO₃) is used as an oxidant. This process takes place under wet and anaerobic (low oxygen) conditions:

NO₃ -> N₂O

This N₂O can however also further react in the complete denitrification process, which can also cause a soil sink in N₂O:

N₂O -> N₂

Figure 12.5: Schematic illustration of the altered nitrogen cycle and nitrous oxide (N₂O) emissions. Figure from Aryal et al. (2022).

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, CH₄ concentrations typically ranged around 320 ppb and were thus substantially lower than during warm (interglacial) periods, when concentrations were around 700 ppb (Figure 12.6). Lower concentrations in glacial period owes to multiple reasons. While ice ages are initially driven by periodic orbital settings, which led to the Earth’s cooling, it also sets off re-enforcing feedbacks of reduced atmospheric concentrations which lead to further cooling. Firstly, a colder ocean stores more carbon as CO₂. Colder temperatures and reduced wetland coverage further lead to the reduction of CH₄ and N₂O emissions from soils. Lastly, the increased formation of ice sheets reduces the planetary albedo, which further cools the climate.

In the the transitions between glacial and interglacial periods 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 12.6: 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. However, rapid changes in atmospheric concentrations occurred also during glacial periods, when CH₄ and N₂O varied rapidly over the course of (mainly northern-hemispheric) climate oscillations (Figure 1.3) associated with variations in the meridional overturning circulation (Arneth et al. 2010).

12.7 Atmospheric sinks of CH₄ and N₂O

As illustrated in Figure 12.2, methane is primarily removed from the atmosphere through chemical loss in the troposphere via oxidation by hydroxyl radicals (OH^-) at 483-682 Tg CH₄ yr^-1. 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. Therefore, the concentration of OH, particularly in the troposphere, is a dominant determinant of the rate of CH₄ removal.

By contrast, as illustrated in Figure 12.4, nitrous oxide is removed in the stratosphere through two principal processes: photolysis, whereby N₂O molecules are broken down by ultraviolet (UV) radiation, and oxidation by O(¹D) radicals, where reactive oxygen atoms oxidize N₂O. Stratospheric 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. In addition, the slight negative feedback of N₂O lifetime to rising N₂O levels results in a marginally shorter residence time.

Common to sink processes for both CH₄ and N₂O is that, to first order, the sink rates are proportional to the respective atmospheric concentration. The atmospheric concentration of, for example N₂O, proportionally scales the number of N₂O molecules exposed to oxidation. Secondary effects modify this rate, but conceiving the sink flux (descruction rate) to be proportional to concentration yields a good approximation.

12.8 Atmospheric lifetime of CH₄ and N₂O

Atmospheric processes are key sinks of both CH4 and N2O in the Earth system, largely compensating 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: \[ \frac{\mathrm{d}C(t)}{\mathrm{d}t} = E(t) - \frac{1}{\tau} C(t) \;. \tag{12.1}\]

\(C(t)\) is the atmospheric mass of either CH₄ or N₂O at time \(t\), and \(E(t)\) are the respective emissions. The proportionality constant \(\tau\) represents the atmospheric lifetime or turnover time (in years) of the greenhouse gas. See Section 3.2 for a more detailed introduction of the concept of the turnover time and the 1st-order decay model.

Methane (CH₄) is a relatively short-lived GHG with an atmospheric lifetime (\(\tau\)) 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 can be rapidly taken up by land or ocean within a few years. The remaining CO2 in the atmosphere however has lifetime of centuries due to direct removal processes in the atmosphere.

NoteGlobal warming potential

Differences in atmospheric residence times of difference greenhouse gases has implication for their warming effects over time. Greenhouse gas effects are often measured relative to effects of carbon dioxide (CO₂), thus in terms of kilograms of CO₂ equivalent. This results in a dimensionless factor known as the Global Warming Potential (GWP). The GWP metric quantifies the cumulative warming effect (or radiative forcing) 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 at the 20-year time scale, than at longer time scales. This owed to the rapid decay of atmospheric CH₄, while 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-fossil	11.8 ± 1.8	5.7 ± 1.4 ×10⁻⁴	82.5 ± 25.8	29.8 ± 11	10.0 ± 3.8
CH4-non fossil	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
HFC-32	5.4 ± 1.1	1.1 ± 0.2 ×10⁻¹	2693 ± 842	771 ± 292	220 ± 87
HFC-134a	14.0 ± 2.8	1.67 ± 0.32 ×10⁻¹	4144 ± 1160	1526 ± 577	436 ± 173
CFC-11	52.0 ± 10.4	2.91 ± 0.65 ×10⁻¹	8321 ± 2419	6226 ± 2297	2093 ± 865
PFC-14	50,000	9.89 ± 0.19 ×10⁻²	5301 ± 1395	7380 ± 2430	10,587 ± 3692

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 removed (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 continues 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) with 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 12.7. 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 12.7: 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 3.8 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, as well as current atmospheric concentration of 1.94 ppm and 0.34. What is the concentration of both gases in the atmosphere after 20, 50, 100 years?
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?

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