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The RF patterns of short-lived climate drivers with inhomogeneous source distributions, such as aerosols, tropospheric ozone, contrails, and land cover change, are leading examples of highly inhomogeneous forcings.

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Spatial and temporal variability in aerosol and aerosol precursor emissions is enhanced by in-atmosphere aerosol formation and chemical transformations, and by aerosol removal in precipitation and surface deposition. The RF of a uniform CO 2 distribution, for example, depends on latitude and cloud cover. Forcing feedbacks in response to spatially variable forcings also have variable geographic and temporal patterns.

Quantifying the relationship between spatial RF patterns and regional and global climate responses in the industrial era is difficult because it requires distinguishing forcing responses from the inherent internal variability of the climate system, which acts on a range of time scales. The ability to test the accuracy of modeled responses to forcing patterns is limited by the sparsity of long-term observational records of regional climate variables.

Chapter 2: Physical Drivers of Climate Change

As a result, there is generally very low confidence in our understanding of the qualitative and quantitative forcing—response relationships at the regional scale. However, there is medium to high confidence in other features, such as aerosol effects altering the location of the Inter Tropical Convergence Zone ITCZ and the positive feedback to reductions of snow and ice and albedo changes at high latitudes.

This radiative feedback, defined as the Planck feedback, only partially offsets the positive RF while triggering other feedbacks that affect radiative balance. These feedback values remain largely unchanged between recent IPCC assessments. A decrease in cloudiness has the opposite effects. Clouds have a relatively larger shortwave effect when they form over dark surfaces for example, oceans than over higher albedo surfaces, such as sea ice and deserts.

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For low-altitude, thick clouds for example, stratus and stratocumulus the shortwave radiative effect dominates, so they cause a net cooling. For high-altitude, thin clouds for example, cirrus the longwave effect dominates, so they cause a net warming e. Therefore, an increase in low clouds is a negative feedback to RF, while an increase in high clouds is a positive feedback. The potential magnitude of cloud feedbacks is large compared with global RF see Section 2.

The value of the shortwave cloud feedback shows a significant sensitivity to computation methodology. Snow and ice are highly reflective to solar radiation relative to land surfaces and the ocean. The losses create the snow—albedo feedback because subsequent increases in absorbed solar radiation lead to further warming as well as changes in turbulent heat fluxes at the surface. For ice sheets for example, on Antarctica and Greenland—see Ch. Specifically, since continental ice shelves limit the discharge rates of ice sheets into the ocean; any melting of the ice shelves accelerates the discharge rate, creating a positive feedback on the ice-stream flow rate and total mass loss e.

Warming oceans also lead to accelerated melting of basal ice ice at the base of a glacier or ice sheet and subsequent ice-sheet loss e. Feedbacks related to ice sheet dynamics occur on longer time scales than other feedbacks—many centuries or longer. Significant ice-sheet melt can also lead to changes in freshwater input to the oceans, which in turn can affect ocean temperatures and circulation, ocean—atmosphere heat exchange and moisture fluxes, and atmospheric circulation. The complete contribution of ice-sheet feedbacks on time scales of millennia are not generally included in CMIP5 climate simulations.

These slow feedbacks are also not thought to change in proportion to global mean surface temperature change, implying that the apparent climate sensitivity changes with time, making it difficult to fully understand climate sensitivity considering only the industrial age.

This slow response increases the likelihood for tipping points, as discussed further in Chapter Potential Surprises. The surface-albedo feedback is an important influence on interannual variations in sea ice as well as on long-term climate change. While there is a significant range in estimates of the snow-albedo feedback, it is assessed as positive, 84 , , with a best estimate of 0. Within the cryosphere, the surface-albedo feedback is most effective in polar regions; 94 , there is also evidence that polar surface-albedo feedbacks might influence the tropical climate as well.

Changes in sea ice can also influence arctic cloudiness. Recent work indicates that arctic clouds have responded to sea ice loss in fall but not summer. Climate change alters the atmospheric abundance and distribution of some radiatively active species by changing natural emissions, atmospheric photochemical reaction rates, atmospheric lifetimes, transport patterns, or deposition rates. These changes in turn alter the associated ERFs, forming a feedback.

Important examples include climate-driven changes in temperature and precipitation that affect 1 natural sources of NO x from soils and lightning and VOC sources from vegetation, all of which affect ozone abundances; , , 2 regional aridity, which influences surface dust sources as well as susceptibility to wildfires; and 3 surface winds, which control the emission of dust from the land surface and the emissions of sea salt and dimethyl sulfide—a natural precursor to sulfate aerosol—from the ocean surface.

Climate-driven ecosystem changes that alter the carbon cycle potentially impact atmospheric CO 2 and CH 4 abundances Section 2. Atmospheric aerosols affect clouds and precipitation rates, which in turn alter aerosol removal rates, lifetimes, and atmospheric abundances. Longwave radiative feedbacks and climate-driven circulation changes also alter stratospheric ozone abundance. While understanding of key processes is improving, atmospheric composition feedbacks are absent or limited in many global climate modeling studies used to project future climate, though this is rapidly changing.


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The cycling of carbon through the climate system is an important long-term climate feedback that affects atmospheric CO 2 concentrations. The global mean atmospheric CO 2 concentration is determined by emissions from burning fossil fuels, wildfires, and permafrost thaw balanced against CO 2 uptake by the oceans and terrestrial biosphere Figures 2. The capacity of the land to continue uptake of CO 2 is uncertain and depends on land-use management and on responses of the biosphere to climate change see Ch.

Altered uptake rates affect atmospheric CO 2 abundance, forcing, and rates of climate change. Such changes are expected to evolve on the decadal and longer time scale, though abrupt changes are possible. Significant uncertainty exists in quantification of carbon-cycle feedbacks, with large differences in the assumed characteristics of the land carbon-cycle processes in current models.

Ocean carbon-cycle changes in future climate scenarios are also highly uncertain. Both of these contribute significant uncertainty to longer-term century-scale climate projections. However, this effect is variable; sometimes plants acclimate so that higher CO 2 concentrations no longer enhance growth e. In addition, CO 2 fertilization is often offset by other factors limiting plant growth, such as water and or nutrient availability and temperature and incoming solar radiation that can be modified by changes in vegetation structure.

With sufficient disturbance, it has been argued that forests could, on net, turn into a source rather than a sink of CO 2. Climate-induced changes in the horizontal for example, landscape to biome and vertical soils to canopy structure of terrestrial ecosystems also alter the physical surface roughness and albedo, as well as biogeochemical carbon and nitrogen cycles and biophysical evapotranspiration and water demand. Combined, these responses constitute climate feedbacks by altering surface albedo and atmospheric GHG abundances.

Drivers of these changes in terrestrial ecosystems include changes in the biophysical growing season, altered seasonality, wildfire patterns, and multiple additional interacting factors Ch.


Accurate determination of future CO 2 stabilization scenarios depends on accounting for the significant role that the land biosphere plays in the global carbon cycle and feedbacks between climate change and the terrestrial carbon cycle. Recent advances in ESMs are beginning to account for other important factors such as nutrient limitations. The majority of the ESMs 7 out of 11 simulated a CO 2 concentration larger by 44 ppm on average than their equivalent non-interactive carbon cycle counterpart. The inclusion of carbon-cycle feedbacks does not alter the lower-end bound on climate sensitivity, but, in most climate models, inclusion pushes the upper bound higher.

The ocean plays a significant role in climate change by playing a critical role in controlling the amount of GHGs including CO 2 , water vapor, and N 2 O and heat in the atmosphere Figure 2. To date most of the net energy increase in the climate system from anthropogenic RF is in the form of ocean heat see Box 3. Marine ecosystems take up CO 2 from the atmosphere in the same way that plants do on land.

Since the ocean is an important carbon sink, climate-driven changes in NPP represent an important feedback because they potentially change atmospheric CO 2 abundance and forcing. There are multiple links between RF-driven changes in climate, physical changes to the ocean, and feedbacks to ocean carbon and heat uptake.

Changes in ocean temperature, circulation, and stratification driven by climate change alter phytoplankton NPP.

Absorption of CO 2 by the ocean also increases its acidity, which can also affect NPP and therefore the carbon sink see Ch. In addition to being an important carbon sink, the ocean dominates the hydrological cycle, since most surface evaporation and rainfall occur over the ocean. Climate warming from radiative forcing also is associated with intensification of the water cycle Ch.

Over decadal time scales the surface ocean salinity has increased in areas of high salinity, such as the subtropical gyres, and decreased in areas of low salinity, such as the Warm Pool region see Ch. Increased stratification inhibits surface mixing, high-latitude convection, and deep-water formation, thereby potentially weakening ocean circulations, in particular the Atlantic Meridional Overturning Circulation AMOC see also Ch.

Observational evidence is mixed regarding whether the AMOC has slowed over the past decades to century see Sect.

Future projections show that the strength of AMOC may significantly decrease as the ocean warms and freshens and as upwelling in the Southern Ocean weakens due to the storm track moving poleward see also Ch. Increased ocean temperatures also accelerate ice sheet melt, particularly for the Antarctic Ice Sheet where basal sea ice melting is important relative to surface melting due to colder surface temperatures.