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TS.5.1 Introduction
Projections of changes in the climate system are made using a hierarchy of climate models ranging from simple climate models, to models of intermediate complexity, to comprehensive climate models, and Earth System Models. These models simulate changes based on a set of scenarios of anthropogenic forcings. A new set of scenarios, the Representative Concentration Pathways (RCPs), was used for the new climate model simulations carried out under the framework of the Coupled Model Intercomparison Project Phase 5 (CMIP5) of the World Climate Research Programme. A large number of comprehensive climate models and Earth System Models have participated in CMIP5, whose results form the core of the climate system projections.
This section summarizes the assessment of these climate change projections. First, future forcing and scenarios are presented. The following subsections then address various aspects of projections of global and regional climate change, including near-term (up to about mid-century) and long-term (end of the 21st century) projections in the atmosphere, ocean and cryosphere; projections of carbon and other biogeochemical cycles; projections in sea level change; and finally changes to climate phenomena and other aspects of regional climate over the 21st century. Projected changes are given relative to the 1986–2005 average unless indicated otherwise. Projections of climate change on longer term and information on climate stabilization and targets are provided in TFE.8. Methods to counter climate change, termed geoengineering, have been proposed and an overview is provided in Box TS.7. {11.3, 12.3–12.5, 13.5–13.7, 14.1–14.6, Annex I}
TS.5.2 Future Forcing and Scenarios
In this assessment report a series of new Representative Concentration Pathways (RCP) are used that largely replace the SRES scenarios (see Box TS.6; see Annex II for Climate System Scenario Tables). They produce a range of responses from ongoing warming, to approximately stabilized forcing, to a stringent mitigation scenario (RCP2.6) that stabilizes and then slowly reduces the radiative forcing after mid-21st century. In contrast to the AR4, the climate change from the RCP scenarios in the AR5 is framed as a combination of adaptation and mitigation. Mitigation actions starting now in the various RCP scenarios do not produce discernibly different climate change outcomes for the next 30 years or so, while long-term climate change after mid-century is appreciably different across the RCPs. {Box 1.1} The range in anthropogenic aerosol emissions across all scenarios has a larger impact on near-term climate projections than the corresponding range in long-lived greenhouse gases, particularly on regional scales and for hydrological cycle variables. The RCP scenarios do not span the range of future aerosol emissions found in the SRES and alternative scenarios (see Box TS.6). {11.3.1, 11.3.6}
If rapid reductions in sulphate aerosol are undertaken for improving air quality or as part of decreasing fossil-fuel CO2 emissions, then there is medium confidence that this could lead to rapid near-term warming. There is robust evidence that accompanying controls on methane (CH4) emissions would offset some of this sulphate-induced warming, although the cooling from methane mitigation will emerge more slowly than the warming from sulphate mitigation due to the different timescales over which atmospheric concentrations of these substances decrease in response to decreases in emissions. While removal of black carbon aerosol could also counter warming associated with sulphate removal, uncertainties are too large to constrain the net sign of the global temperature response to black carbon emission reductions, which depends on reduction of co-emitted (reflective) aerosols and on aerosol indirect effects. {11.3.6} Including uncertainties in projecting the chemically reactive greenhouse gases methane (CH4) and nitrous oxide (N2O) from RCP emissions gives a range in abundance pathways that is likely 30% larger than the range in RCP concentrations used to force the CMIP5 climate models. Including uncertainties in emission estimates from agricultural, forest, and land-use sources, in atmospheric lifetimes, and in chemical feedbacks, results in a much wider range of abundances for N2O, CH4, and HFCs and their radiative forcing.
In the case of CH4 it likely extends the range up to 500 ppb above RCP8.5 and 270 ppb below RCP2.6 through to 2100, with smaller ranges in the near term. {11.3.5} There is low confidence in projections of natural forcing. Major volcanic eruptions cause a negative radiative forcing up to several W m–2, with a typical lifetime of one year, but the possible occurrence and timing of future eruptions is unknown. Except for the 11-year solar cycle, changes in the total solar irradiance are uncertain. Except where explicitly indicated, future volcanic eruptions and changes in total solar irradiance additional to a repeating 11 year solar cycle are not included in the projections of near- and long-term climate assessed. {8, 11.3.1}
TS.5.3 Quantification of Climate System Response
Estimates of the Equilibrium Climate Sensitivity (ECS) based on observed climate change, climate models and feedback analysis, as well as paleoclimate evidence indicate that ECS is positive, likely in the range 1.5°C to 4.5°C with high confidence, extremely unlikely less than 1°C (high confidence) and very unlikely greater than 6°C (medium confidence). Earth system sensitivity over millennia timescales including long- term feedbacks not typically included in models could be significantly higher than ECS (see TFE.6 for further details). {5.3.1, 10.8; Box 12.2}
With high confidence the transient climate response (TCR) is positive, likely in the range 1°C to 2.5oC and extremely unlikely greater than 3°C, based on observed climate change and climate models (see TFE.6 for further details). {10.8; Box 12.2} The ratio of global mean surface temperature change to total cumulative anthropogenic carbon emissions is relatively constant and independent of the scenario, but is model dependent, as it is a function of the model cumulative airborne fraction of carbon and the transient climate response. For any given temperature target, higher emissions in earlier decades therefore imply lower emissions by about the same amount later on. The transient climate response to cumulative carbon emission (TCRE) is likely between 0.8°C to 2.5°C per 1000 PgC (high confidence), for cumulative carbon emissions less than about 2000PgC until the time at which temperatures peak (see TFE.8 for further details). {10.8, 12.5.4; Box 12.2}
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Box TS.6: The New RCP Scenarios and CMIP5 Models
Future anthropogenic emissions of greenhouse gases (GHG), aerosol particles and other forcing agents such as land use change are dependent on socio-economic factors, and may be affected by global geopolitical agreements to control those emissions to achieve mitigation. AR4 made extensive use of the SRES scenarios that do not include additional climate initiatives, which means that no scenarios were included that explicitly assume implementation of the United Nations Framework Convention on Climate Change (UNFCCC) or the emissions targets of the Kyoto Protocol. However, GHG emissions are directly affected by non-climate change policies designed for a wide range of other purposes. The SRES scenarios were developed using a sequential approach, i.e., socio-economic factors fed into emissions scenarios, which were then used in simple climate models to determine concentrations of greenhouse gases, and other agents required to drive the more complex atmosphere-ocean global climate models (AOGCMs). In this report, outcomes of climate simulations that use new scenarios (some of which include implied policy actions to achieve mitigation) referred to as “Representative Concentration Pathways” (RCPs) are assessed. These RCPs represent a larger set of mitigation scenarios and were selected to have different targets in terms of radiative forcing at 2100 (about 2.6, 4.5, 6.0 and 8.5 W m–2; see Figure TS.15). The scenarios should be considered plausible and illustrative, and do not have probabilities attached to them. {12.3.1; Box 1.1}
The RCPs were developed using Integrated Assessment Models (IAMs) that typically include economic, demographic, energy, and simple climate components. The emission scenarios they produce are then run through a simple model to produce time series of greenhouse gas concentrations that can be run in AOGCMs. The emission time series from the RCPs can then be used directly in Earth System Models (ESMs) that include interactive biogeochemistry (at least a land and ocean carbon cycle).
The CMIP5 multi-model experiment (coordinated through the World Climate Research Programme) presents an unprecedented level of information on which to base assessments of climate variability and change. CMIP5 includes new ESMs in addition to AOGCMs, new model experiments, and more diagnostic output. CMIP5 is much more comprehensive than the preceding CMIP3 multi-model experiment that was available at the time of the IPCC AR4. CMIP5 has more than twice as many models, many more experiments (that also include experiments to address understanding of the responses in the future climate change scenario runs), and nearly 2 1015 bytes of data (as compared to over 30 1012 bytes of data in CMIP3). A larger number of forcing agents are treated more completely in the CMIP5 models, with respect to aerosols and land use particularly. Black carbon aerosol is now a commonly included forcing agent. Considering CO2, both ‘concentrations-driven’ projections and ‘emissions-driven’ projections are assessed from CMIP5. These allow quantification of the physical response uncertainties as well as climate-carbon cycle interactions. {1.5.2} The assessment of the mean values and ranges of global mean temperature changes in AR4 would not have been substantially different if the CMIP5 models had been used in that report. The differences in global temperature projections can largely be attributed to the different scenarios. The global mean temperature response simulated by CMIP3 and CMIP5 models is very similar, both in the mean and the model range, transiently and in equilibrium. The range of temperature change across all scenarios is wider because the RCPs include a strong mitigation scenario (RCP2.6) that had no equivalent among the SRES scenarios used in CMIP3. For each scenario, the 5–95% range of the CMIP5 projections is obtained by approximating the CMIP5 distributions by a normal distribution with same mean and standard deviation and assessed as being “likely” for projections of global temperature change for the end of the 21st century. Probabilistic projections with simpler models calibrated to span the range of equilibrium climate sensitivity assessed by the AR4 provide uncertainty ranges that are consistent with those from CMIP5. In AR4 the uncertainties in global temperature projections were found to be approximately constant when expressed as a fraction of the model mean warming (constant fractional uncertainty). For the higher RCPs, the uncertainty is now estimated to be smaller than with the AR4 method for long-term climate change, because the carbon cycle climate feedbacks are not relevant for the concentration driven RCP projections (in contrast, the assessed projection uncertainties of global temperature in AR4 did account of carbon cycle climate feedbacks, even though these were not part of the CMIP3 models). When forced with RCP8.5, CO2 emissions, as opposed to the RCP8.5 CO2 concentrations, CMIP5 Earth System Models (ESMs) with interactive carbon cycle simulate, on average, a 50 (–140 to +210) ppm (CMIP5 model spread) larger atmospheric CO2 concentration and 0.2°C larger global surface temperature increase by 2100. For the low RCPs the fractional uncertainty is larger because internal variability and non-CO2 forcings make a larger relative contribution to the total uncertainty. {12.4.1, 12.4.9} There is overall consistency between the projections of temperature and precipitation based on CMIP3 and CMIP5, both for large-scale patterns and magnitudes of change (Box TS.6, Figure 1). Model agreement and confidence in projections depends on the variable and on spatial and temporal averaging, with better agreement for larger scales. Confidence is higher for temperature than for those quantities related to the water cycle or atmospheric circulation. Improved methods to quantify and display model robustness have been developed to indicate where lack of agreement across models on local trends is a result of internal variability, rather than models actually disagreeing on their forced response. Understanding of the sources and means of characterizing uncertainties in long-term large scale projections of climate change has not changed significantly since AR4, but new experiments and studies have continued to work towards a more complete and rigorous characterization. {9.7.3, 12.2, 12.4.1, 12.4.4, 12.4.5, 12.4.9; Box 12.1}
The well-established stability of geographical patterns of temperature and precipitation change during a transient experiment remains valid in the CMIP5 models (see Box TS.6, Figure 1). Patterns are similar over time and across scenarios and to first order can be scaled by the global mean temperature change. There remain limitations to the validity of this technique when it is applied to strong mitigation scenarios, to scenarios where localized forcing (e.g., aerosols) are significant and vary in time and for variables other than average seasonal mean temperature and precipitation {12.4.2}.
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TFE.6: Climate Sensitivity and Feedbacks
The description of climate change as a response to a forcing that is amplified by feedbacks goes back many decades. The concepts of radiative forcing and climate feedbacks continue to be refined, and limitations are now better understood; for instance, feedbacks may be much faster than the surface warming, feedbacks depend on the type of forcing agent (e.g., greenhouse gas vs. solar forcing), or may have intrinsic timescales (associated mainly with vegetation change and ice sheets) of several centuries to millennia. The analysis of physical feedbacks in models and from observations remains a powerful framework that provides constraints on transient future warming for different scenarios, on climate sensitivity and, combined with estimates of carbon cycle feedbacks (see TFE.5), determines the greenhouse gas emissions that are compatible with climate stabilization or targets (see TFE.8). {7.1, 9.7.2, 12.5.3; Box 12.2} The water vapour/lapse rate, albedo and cloud feedbacks are the principal determinants of equilibrium climate sensitivity (ECS, the equilibrium change in annual mean global surface temperature following a doubling of the atmospheric CO2 concentration). All of these feedbacks are assessed to be positive, but with different levels of likelihood assigned ranging from likely to extremely likely. Therefore, there is very high confidence that the net feedback is strongly positive and the black body response of the climate to a forcing will therefore be amplified. Cloud feedbacks continue to be the largest uncertainty. The net feedback from water vapour and lapse rate changes together is extremely likely positive and approximately doubles the black body response. The mean value and spread of these two processes in climate models are essentially unchanged from AR4, but are now supported by stronger observational evidence and better process understanding of what determines relative humidity distributions.. Clouds respond to climate forcing mechanisms in multiple ways and individual cloud feedbacks can be positive or negative. Key issues include the representation of both deep and shallow cumulus convection, microphysical processes in ice clouds, and partial cloudiness that results from small-scale variations of cloud-producing and cloud-dissipating processes. New approaches to diagnosing cloud feedback in GCMs have clarified robust cloud responses, while continuing to implicate low cloud cover as the most important source of intermodel spread in simulated cloud feedbacks. The net radiative feedback due to all cloud types is likely positive. This conclusion is reached by considering a plausible range for unknown contributions by processes yet to be accounted for, in addition to those occurring in current climate models. Observations alone do not currently provide a robust, direct constraint, but multiple lines of evidence now indicate positive feedback contributions from changes in both the height of high clouds and the horizontal distribution of clouds. The additional feedback from low cloud amount is also positive in most climate models, but that result is not well understood, nor effectively constrained by observations, so confidence in it is low. {7.2.4, 7.2.5, 7.2.6}
The representation of aerosol-cloud processes in climate models continues to be a challenge. Aerosol and cloud variability at scales significantly smaller than those resolved in climate models, and the subtle responses of clouds to aerosol at those scales, mean that, for the foreseeable future, climate models will continue to rely on parameterizations of aerosol-cloud interactions or other methods that represent subgrid variability. This implies large uncertainties for estimates of the forcings associated with aerosol-cloud interactions. {7.4, 7.5.3, 7.5.4} Equilibrium climate sensitivity (ECS) and Transient Climate Response (TCR) are useful metrics summarising the global climate system's temperature response to an externally-imposed radiative forcing. ECS is defined as the equilibrium change in annual mean global mean surface temperature following a doubling of the atmospheric CO2 concentration (see Glossary), while TCR is defined as the annual mean global mean surface temperature change at the time of CO2 doubling following a linear increase in CO2 forcing over a period of 70 years (see Glossary). Both metrics have a broader application than these definitions imply: ECS determines the eventual warming in response to stabilisation of atmospheric composition on multi-century timescales, while TCR determines the warming expected at a given time following any steady increase in forcing over a 50- to 100-year timescale. {Box 12.2, 12.5.3} ECS and TCR can be estimated from various lines of evidence (see TFE.6, Figures 1 and 2). The estimates can be based on the values of ECS and TCR diagnosed from climate models, or they can be constrained by analysis of feedbacks in climate models, patterns of mean climate and variability in models compared to observations, temperature fluctuations as reconstructed from paleoclimate archives, observed and modelled short term perturbations of the energy balance like those caused by volcanic eruptions, and the observed surface and ocean temperature trends since preindustrial. For many applications, the limitations of the forcing-feedback analysis framework and the dependence of feedbacks on timescales and the climate state must be kept in mind. {5.3.1, 5.3.3, 9.7.1, 9.7.2, 9.7.3, 10.8.1, 10.8.2, 12.5.3; Box 5.2; Table 9.5}
Newer studies of constraints on ECS are based on the observed warming since preindustrial, analysed using simple and intermediate complexity models, improved statistical methods, and several different and newer datasets. Together with results from feedback analysis and paleoclimate constraints these studies show ECS is likely between 1.5°C to 4.5°C (medium confidence) and extremely unlikely less than 1.0°C. {5.3.1, 5.3.3, 10.8.2; Box 5.2, 12.2} Estimates based on AOGCMs and feedback analysis indicate a range of 2 to 4.5°C, with the CMIP5 model mean at 3.2°C, similar to CMIP3. High climate sensitivities are found in some perturbed parameter ensembles models, but recent comparisons of perturbed-physics ensembles against the observed climate find that models with ECS values in the range 3 to 4°C show the smallest errors for many fields. Relationships between climatological quantities and climate sensitivity are often found within a specific perturbed parameter ensemble model but in many cases the relationship is not robust across perturbed parameter ensembles models from different models or in CMIP3/5. The assessed literature suggests that the range of climate sensitivities and transient responses covered by CMIP3/5 cannot be narrowed significantly by constraining the models with observations of the mean climate and variability. Studies based on perturbed parameter ensembles models and CMIP3 support the conclusion that a credible representation of the mean climate and variability is very difficult to achieve with ECSs below 2°C. {9.2.2, 9.7.3; Box 12.2}
New estimates of ECS based on reconstructions and simulations of the Last Glacial Maximum (21,000 years to 19,000 years ago) show that values below 1°C as well as above 6°C are very unlikely. In some models climate sensitivity differs between warm and cold climates because of differences in the representation of cloud feedbacks. Estimates of an Earth System sensitivity including slow feedbacks (e.g., ice sheets or vegetation) are even more difficult to relate to climate sensitivity of the current climate state. The main limitations of ECS estimates from paleoclimate states are uncertainties in proxy data, spatial coverage of the data, uncertainties in some forcings, and structural limitations in models used in model-data comparisons. {5.3, 10.8.2, 12.5.3}
Bayesian methods to constrain ECS or TCR are sensitive to the assumed prior distributions. They can in principle yields narrower estimates by combining constraints from the observed warming trend, volcanic eruptions, model climatology, and paleoclimate, and that has been done in some studies, but there is no consensus on how this should be done robustly. This approach is sensitive to the assumptions regarding the independence of the various lines of evidence, the possibility of shared biases in models or feedback estimates, and the assumption that each individual line of evidence is unbiased. The combination of different estimates in this assessment is based on expert judgment. {10.8.2; Box 12.2}
Based on the combined evidence from observed climate change including the observed 20th century warming, climate models, feedback analysis and paleoclimate, as discussed above, equilibrium climate sensitivity (ECS) is likely in the range 1.5 to 4.5°C with high confidence. ECS is positive, extremely unlikely less than 1°C (high confidence), and very unlikely greater than 6°C (medium confidence). The tails of the ECS distribution are now better understood. Multiple lines of evidence provide high confidence that an ECS value less than 1°C is extremely unlikely. The upper limit of the likely range is unchanged compared to AR4. The lower limit of the likely range of 1.5°C is less than the lower limit of 2°C in AR4. This change reflects the evidence from new studies of observed temperature change, using the extended records in atmosphere and ocean. These studies suggest a best fit to the observed surface and ocean warming for ECS values in the lower part of the likely range. Note that these studies are not purely observational, because they require an estimate of the response to radiative forcing from models. In addition, the uncertainty in ocean heat uptake remains substantial. Accounting for short-term variability in simple models remains challenging, and it is important not to give undue weight to any short time period which might be strongly affected by internal variability. On the other hand, AOGCMs show very good agreement with observed climatology with ECS values in the upper part of the 1.5-4.5°C range, but the simulation of key feedbacks like clouds remains challenging in those models. The estimates from the observed warming, paleoclimate, and from climate models are consistent within their uncertainties, each is supported by many studies and multiple datasets, and in combination they provide high confidence for the assessed likely range. Even though this assessed range is similar to previous reports, confidence today is much higher as a result of high quality and longer observational records with a clearer anthropogenic signal, better process understanding, more and better understood evidence from paleoclimate reconstructions, and better climate models with higher resolution that capture many more processes more realistically. All these lines of evidence individually support the assessed likely range of 1.5 to 4.5°C. {3.2, 9.7.3, 10.8; Box 9.2, 13.1}
On timescales of many centuries and longer, additional feedbacks with their own intrinsic timescales (e.g., vegetation, ice sheets) may become important but are not usually modelled in AOGCMs. The resulting equilibrium temperature response to a doubling of CO2 on millennial timescales or Earth System Sensitivity is less well constrained but likely to be larger than ECS, implying that lower atmospheric CO2 concentrations are compatible with limiting warming to below a given temperature level. These slow feedbacks are less likely to be proportional to global mean temperature change, implying that Earth System Sensitivity changes over time. Estimates of Earth System Sensitivity are also difficult to relate to climate sensitivity of the current climate state. {5.3.3, 10.8.2, 12.5.3}
For scenarios of increasing radiative forcing, TCR is a more informative indicator of future climate change than ECS. This assessment concludes with high confidence that the transient climate response (TCR) is likely in the range 1°C to 2.5°C, close to the estimated 5–95% range of CMIP5 (1.2°C to 2.4°C), is positive and extremely unlikely greater than 3°C. As with the ECS, this is an expert-assessed range, supported by several different and partly independent lines of evidence, each based on multiple studies, models and datasets. TCR is estimated from the observed global changes in surface temperature, ocean heat uptake and radiative forcing including detection/attribution studies identifying the response patterns to increasing greenhouse gas concentrations, and the results of CMIP3 and CMIP5. Estimating TCR suffers from fewer difficulties in terms of state- or time-dependent feedbacks, and is less affected by uncertainty as to how much energy is taken up by the ocean. Unlike ECS, the ranges of TCR estimated from the observed warming and from AOGCMs agree well, increasing our confidence in the assessment of uncertainties in projections over the 21st century. The assessed ranges of ECS and TCR are largely consistent with the observed warming, the estimated forcing, and the projected future warming. {9.7.1, 10.8.1, 12.5.3; Table 9.5}
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TS.5.6 Long-Term Projections of Carbon and Other Biogeochemical Cycles Projections of the global carbon cycle to 2100 using the CMIP5 Earth System Models (ESMs) represent a wider range of complex interactions between the carbon cycle and the physical climate system.
With very high confidence, ocean carbon uptake of anthropogenic CO2 will continue under all four Representative Concentration Pathways through to 2100, with higher uptake in higher concentration pathways. The future evolution of the land carbon uptake is much more uncertain.
A majority of CMIP5 ESMs project a continued net carbon uptake by land ecosystems through 2100. Yet, a minority of models simulate a net CO2 source to the atmosphere by 2100 due to the combined effect of climate change and land use change. In view of the large spread of model results and incomplete process representation, there is low confidence on the magnitude of modelled future land carbon changes. {6.4.3}
There is high confidence that climate change will partially offset increases in global land and ocean carbon sinks caused by rising atmospheric CO2. Yet, there are regional differences among CMIP5 ESMs, in the response of ocean and land CO2 fluxes to climate. There is high agreement between models that tropical ecosystems will store less carbon in a warmer climate. There is medium agreement between the CMIP5 ESMs that at high latitudes warming will increase land carbon storage, although none of these models accounts for decomposition of carbon in permafrost, which may offset increased land carbon storage. There is high confidence that reductions in permafrost extent due to warming will cause thawing of some currently frozen carbon. However, there is low confidence on the magnitude of carbon losses through CO2 and CH4 emissions to the atmosphere with a range from 50 to more than 250 PgC between 2000 and 2100 for RCP8.5. {6.4.2, 6.4.3}
The loss of carbon from frozen soils constitutes a positive radiative feedback that is missing in current coupled ESM projections. There is high agreement between CMIP5 ESMs that ocean warming and circulation changes will reduce the rate of ocean carbon uptake in the Southern Ocean and North Atlantic, but that carbon uptake will nevertheless persist in those regions. {6.4.2}
It is very likely, based on new experimental results and modelling, that nutrient shortage will limit the effect of rising atmospheric CO2 on future land carbon sinks, for the four RCP scenarios. There is high confidence that low nitrogen availability will limit carbon storage on land, even when considering anthropogenic nitrogen deposition. The role of phosphorus limitation is more uncertain. {6.4.6}
For the ESMs simulations driven by CO2 concentrations, representation of the land and ocean carbon cycle allows quantification of the fossil fuel emissions compatible with the RCP scenarios. Between 2012 and 2100, ESM results imply cumulative compatible fossil fuel emissions of 270 [140 to 410] PgC for RCP2.6, 780 [595 to 1005] PgC for RCP4.5, 1060 [840 to 1250] PgC for RCP6.0, and 1685 [1415 to 1910] PgC for RCP8.5 (CMIP5 model spread) (Figure TS.19). For RCP2.6, the models project an average 50% (range 14– 96%) emission reduction by 2050 relative to 1990 levels. It is about as likely as not that sustained globally negative emissions will be required to achieve the reductions in atmospheric CO2 in RCP2.6. See also Box TS.7. {6.4.3; Table 6.12}
When forced with RCP8.5 CO2 emissions, as opposed to the RCP8.5 CO2 concentrations, CMIP5 ESMs with interactive carbon cycles simulate, on average, a 50 (–140 to +210) ppm (CMIP5 model spread) larger atmospheric CO2 concentration and a 0.2 (–0.4 to +0.9) °C (CMIP5 model spread) larger global surface temperature increase by 2100. {12.4.8} It is virtually certain that the increased storage of carbon by the ocean will increase acidification in the future, continuing the observed trends of the past decades. Ocean acidification in the surface ocean will follow atmospheric CO2 and it will also increase in the deep ocean as CO2 continues to penetrate the abyss. The CMIP5 models consistently project worldwide increased ocean acidification to 2100 under all RCPs. The corresponding decrease in surface ocean pH by the end of 21st century is 0.065 (0.06 to 0.07) for RCP2.6, 0.145 (0.14 to 0.15) for RCP4.5, 0.203 (0.20 to 0.21) for RCP6.0, and 0.31 (0.30 to 0.32) for RCP8.5 (CMIP5 model spread) (Figure TS.20). Surface waters are projected to become seasonally corrosive to aragonite in parts of the Arctic and in some coastal upwelling systems within a decade, and in parts of the Southern Ocean within 1–3 decades in most scenarios. Aragonite, a less stable form of calcium carbonate, undersaturation becomes widespread in these regions at atmospheric CO2 levels of 500–600 ppm. {6.4.4}
It is very likely that the dissolved oxygen content of the ocean will decrease by 3 to 6% during the 21st century in response to surface warming (see Box 6.5, and Section 6.4.5). CMIP5 models suggest that this decrease in dissolved oxygen will predominantly occur in the subsurface mid-latitude oceans, caused by enhanced stratification, reduced ventilation and warming. However, there is no consensus on the future development of the volume of hypoxic and suboxic waters in the open-ocean because of large uncertainties in potential biogeochemical effects and in the evolution of tropical ocean dynamics. {6.4.5} With very high confidence, the physical, biogeochemical carbon cycle in the ocean and on land will continue to respond to climate change and atmospheric CO2 increases that arise during the 21st century (see TFE.7 and TFE 8).
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TFE.7: Carbon Cycle Perturbation and Uncertainties The natural carbon cycle has been perturbed since the beginning of the Industrial Revolution (circa 1750) by the anthropogenic release of CO2 to the atmosphere, virtually all from fossil fuel combustion and land use change, with a small contribution from cement production. Fossil fuel burning is a process related to energy production. Fossil fuel carbon comes from geological deposits of coal, oil and gas that were buried in the Earth crust for millions of years. Land use change CO2 emissions are related to the conversion of natural ecosystems into managed ecosystems for food, feed and timber production with CO2 being emitted from the burning of plant material or from the decomposition of dead plants and soil organic carbon. For instance when a forest is cleared, the plant material may be released to the atmosphere quickly through burning or over many years as the dead biomass and soil carbon decay on their own. {6.1, 6.3; Table 6.1}
The human caused excess of CO2 in the atmosphere is partly removed from the atmosphere by carbon sinks in land ecosystems and in the ocean, currently leaving about 45% of the CO2 emissions in the atmosphere. Natural carbon sinks are due to physical, biological and chemical processes acting on different time scales. An excess of atmospheric CO2 supports photosynthetic CO2 fixation by plants that is stored as plant biomass or in the soil. The residence times of stored carbon on land depends on the compartments (plant / soil) and composition of the organic carbon, with time horizons varying from days to centuries. The increased storage in terrestrial ecosystems not affected by land use change is likely to be caused by enhanced photosynthesis at higher CO2 levels and N deposition, and changes in climate favoring carbon sinks such as longer growing seasons in mid-to-high latitudes. {6.3, 6.3.1}
The uptake of anthropogenic CO2 by the ocean is primarily a response to increasing CO2 in the atmosphere. Excess atmospheric CO2 absorbed by the surface ocean or transported to the ocean through aquatic systems (e.g., rivers, ground waters) gets buried in coastal sediments or transported to deep waters where it is stored for decades to centuries. The deep ocean carbon can dissolve ocean carbonate sediments to store excess CO2 on time scales of centuries to millennia. Within a thousand years, the remaining atmospheric fraction of the CO2 emissions will be between 15 and 40%, depending on the amount of carbon released (TFE.7, Figure 1). On geological time scales of 10,000 years or longer, additional CO2 is removed very slowly from the atmosphere by rock weathering, pulling the remaining atmospheric CO2 fraction down to 10 to 25% after 10,000 years. {Box 6.1}
The carbon cycle response to future climate and CO2 changes can be viewed as two strong and opposing feedbacks. The concentration-carbon feedback determines changes in storage due to elevated CO2, and the climate-carbon feedback determines changes in carbon storage due to changes in climate. There is high confidence that increased atmospheric CO2 will lead to increased land and ocean carbon uptake but by an uncertain amount. Models agree on the positive sign of land and ocean response to rising CO2 but show only medium and low agreement for the magnitude of ocean and land carbon uptake respectively (TFE.7, Figure 2). Future climate change will decrease land and ocean carbon uptake compared to the case with constant climate (medium confidence). This is further supported by paleoclimate observations and modelling indicating that there is a positive feedback between climate and the carbon cycle on century to millennial time scales. Models agree on the sign, globally negative, of land and ocean response to climate change but show low agreement on the magnitude of this response, especially for the land (TFE.7, Figure 2). A key update since AR4 is the introduction of nutrient dynamics in some land carbon models, in particular the limitations on plant growth imposed by nitrogen availability. There is high confidence that, at the global scale, relative to CMIP5 carbon only ESMs, CMIP5 ESMs including a land nitrogen cycle will reduce the strength of both the concentration-carbon feedback and the climate-carbon feedback of land ecosystems (TFE.7, Figure 2). Inclusion of N-cycle processes increases the spread across the CMIP5 ensemble. The CMIP5 spread in ocean sensitivity to CO2 and climate appears reduced compared to AR4 (TFE.7, Figure 2). {6.2.3, 6.4.2}
With very high confidence, ocean carbon uptake of anthropogenic CO2 emissions will continue under all four Representative Concentration Pathways through to 2100, with higher uptake corresponding to higher concentration pathways. The future evolution of the land carbon uptake is much more uncertain, with a majority of models projecting a continued net carbon uptake under all RCPs, but with some models simulating a net loss of carbon by the land due to the combined effect of climate change and land use change. In view of the large spread of model results and incomplete process representation, there is low confidence on the magnitude of modelled future land carbon changes. [6.4.3, Figure 6.24]
Biogeochemical cycles and feedbacks other than the carbon cycle play an important role in the future of the climate system, although the carbon cycle represents the strongest of these. Changes in the nitrogen cycle, in addition to interactions with CO2 sources and sinks, affect emissions of N2O both on land and from the ocean. The human-caused creation of reactive nitrogen has increased steadily over the last two decades and is dominated by the production of ammonia for fertilizer and industry, with important contributions from legume cultivation and combustion of fossil fuels. {6.3}
Many processes, however, are not yet represented in coupled climate-biogeochemistry models (e.g., other processes involving other biogenic elements such as P, Si, Fe) so their magnitudes have to be estimated in offline or simpler models, which make their quantitative assessment difficult. It is likely that there will be non-linear interactions between many of these processes, but these are not yet well quantified. Therefore any assessment of the future feedbacks between climate and biogeochemical cycles still contains large uncertainty. {6.4}
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TS.5.4 Near-Term Climate Change
Near-term decadal climate prediction provides information not available from existing seasonal to interannual (months to a year or two) predictions or from long-term (mid 21st century and beyond) climate change projections. Prediction efforts on seasonal to interannual timescales require accurate estimates of the initial climate state with less focus extended to changes in external forcing12, while long-term climate projections rely more heavily on estimations of external forcing with little reliance on the initial state of internal variability. Estimates of near-term climate depend on the committed warming (caused by the inertia of the oceans as they respond to historical external forcing) the time evolution of internally-generated climate variability, and the future path of external forcing. Near-term predictions out to about a decade (Figure TS.13) depend more heavily on an accurate depiction of the internally generated climate variability. {11.1, 12, 14}
Further near-term warming from past emissions is unavoidable due to thermal inertia of the oceans. This warming will be increased by ongoing emissions of GHGs over the near term, and the climate observed in the near term will also be strongly influenced by the internally generated variability of the climate system. Previous IPCC Assessments only described climate-change projections wherein the externally forced component of future climate was included but no attempt was made to initialize the internally generated climate variability. Decadal climate predictions, on the other hand, are intended to predict both the externally forced component of future climate change, and the internally generated component. Near-term predictions do not provide detailed information of the evolution of weather. Instead they can provide estimated changes in the time evolution of the statistics of near-term climate. {11.1, 11.2.2; Box 11.1, FAQ 11.1}
Retrospective prediction experiments have been used to assess forecast quality. There is high confidence that the retrospective prediction experiments for forecast periods of up to 10 years exhibit positive skill when verified against observations over large regions of the planet and of the global mean. Observation-based initialization of the forecasts contributes to the skill of predictions of annual mean temperature for the first couple of years and to the skill of predictions of the global-mean surface temperature and the temperature over the North Atlantic, regions of the South Pacific and the tropical Indian Ocean up to 10 years (high confidence) partly due to a correction of the forced response. Probabilistic temperature predictions are statistically reliable (see Section 11.2.3 for definition of reliability) due to the correct representation of global trends, but still unreliable at the regional scale when probabilities are computed from the multi-model ensemble (medium confidence). Predictions initialized over 2000–2005 improve estimates of the recent global-mean temperature hiatus (medium confidence). Predictions of precipitation over continental areas with large forced trends also exhibit positive skill. {11.2.2, 11.2.3; Box 9.2}
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TS.5.4.1. Projected Near-Term Changes in Climate
Projections of near-term climate show small sensitivity to greenhouse gas scenarios compared to model spread, but substantial sensitivity to uncertainties in aerosol emissions, especially on regional scales and for hydrological cycle variables (high confidence). In some regions, the local and regional responses in precipitation and in mean and extreme temperature to land use change will be larger than those due to large- scale greenhouse gases and aerosol forcing (medium confidence). These scenarios presume that there are no major volcanic eruptions and that anthropogenic aerosol emissions are rapidly reduced during the near term.. {11.3.1, 11.3.2, 11.3.6}
TS.5.4.2 Projected Near-Term Changes in Temperature
In the absence of major volcanic eruptions—which would cause significant but temporary cooling—and, assuming no significant future long term changes in solar irradiance, it is likely that the GMST anomaly for the period 2016–2035, relative to the reference period of 1986–2005 will be in the range 0.3°C to 0.7°C (medium confidence). This is based on an assessment of observationally-constrained projections and predictions initialized with observations. This range is consistent with the range obtained by using CMIP5 5– 95% model trends for 2012–2035. It is also consistent with the CMIP5 5–95% range for all four RCP scenarios of 0.36°C to 0.79°C, using the 2006–2012 reference period, after the upper and lower bounds are reduced by 10% to take into account the evidence that some models may be too sensitive to anthropogenic forcing (see Table TS.1 and Figure TS.14). {11.3.6}
There is high confidence that higher concentrations of greenhouse gases and lower amounts of sulphate aerosol lead to greater warming. In the near-term, differences in global mean surface air temperature across RCP scenarios for a single climate model are typically smaller than across climate models for a single RCP scenario. In 2030, the CMIP5 multi-model ensemble mean values for global mean temperature differ by less than 0.3°C between the RCP scenarios, whereas the model spread (defined as the 5–95% range of the decadal means of the models) is around 0.8°C. The inter-scenario spread increases in time and by 2050 is comparable to the model spread. Regionally, the largest differences in surface air temperature between RCP scenarios are found in the Arctic. {11.3.2. 11.3.6}
The projected warming of global mean temperatures implies high confidence that new levels of warming relative to pre-industrial climate will be crossed, particularly under higher greenhouse gas emissions scenarios. Relative to a reference period of 1850–1900, under RCP4.5 or RCP6.0, it is more likely than not that the mean GMST for the period 2016–2035 will be more than 1°C above the mean for 1850–1900, and very unlikely that it will be more than 1.5°C above the 1850–1900 mean (medium confidence). {11.3.6}
A future volcanic eruption similar in size to the 1991 eruption of Mount Pinatubo would cause a rapid drop in global mean surface air temperature of about 0.5°C in the following year, with recovery over the next few years. Larger eruptions, or several eruptions occurring close together in time, would lead to larger and more persistent effects. {11.3.6} Possible future changes in solar irradiance could influence the rate at which global mean surface air temperature increases, but there is high confidence that this influence will be small in comparison to the influence of increasing concentrations of greenhouse gases in the atmosphere. {11.3.6}
The spatial patterns of near-term warming projected by the CMIP5 models following the RCP scenarios (see Figure TS.15) are broadly consistent with the AR4. It is very likely that anthropogenic warming of surface air temperature over the next few decades will proceed more rapidly over land areas than over oceans, and it is very likely that the anthropogenic warming over the Arctic in winter will be greater than the global mean warming, consistent with the AR4. Relative to background levels of internally generated variability there is high confidence that the anthropogenic warming relative to the reference period is expected to be larger in the tropics and subtropics than in mid-latitudes. {11.3.2}
It is likely that in the next decades the frequency of warm days and warm nights will increase in most regions, while the frequency of cold days and cold nights will decrease. Models also project increases in the duration, intensity and spatial extent of heat-waves and warm spells for the near term. These changes may proceed at a different rate than the mean warming. For example, several studies project that European high-percentile summer temperatures are projected to warm faster than mean temperatures (see also TFE.9). {11.3.2}
TS.5.4.3 Projected Near-Term Changes in the Water Cycle
Zonal mean precipitation will very likely increase in high and some of the mid latitudes, and will more likely than not decrease in the subtropics. At more regional scales precipitation changes may be dominated by a combination of natural internal variability, volcanic forcing and anthropogenic aerosol effects. {11.3.2}
Over the next few decades increases in near-surface specific humidity are very likely. It is likely that there will be increases in evaporation in many regions. There is low confidence in projected changes in soil moisture and surface run off. {11.3.2}
In the near term, it is likely that the frequency and intensity of heavy precipitation events will increase over land. These changes are primarily driven by increases in atmospheric water vapor content, but also affected by changes in atmospheric circulation. The impact of anthropogenic forcing at regional scales is less obvious, as regional-scale changes are strongly affected by natural variability and also depend upon the course of future aerosol emissions, volcanic forcing and land use changes (see also TFE.9). {11.3.2}
TS.5.4.4 Projected Near-Term Changes in Atmospheric Circulation
Internally generated climate variability and multiple radiative forcing agents (e.g., volcanoes, greenhouse gases, ozone and anthropogenic aerosols) will all contribute to near-term changes in the atmospheric circulation. For example, it is likely that the annual-mean Hadley Circulation and the Southern Hemisphere mid-latitude westerlies will shift poleward, while it is likely that the projected recovery of stratospheric ozone and increases in greenhouse gas concentrations will have counteracting impacts on the width of the Hadley Circulation and the meridional position of the Southern Hemisphere storm track. Therefore it is unlikely that they will continue to expand poleward as rapidly as in recent decades. {11.3.2}
There is low confidence in near-term projections of the position and strength of Northern Hemisphere storm tracks. Natural variations are larger than the projected impact of greenhouse gases in the near-term. {11.3.2}
There is low confidence in basin-scale projections of changes in intensity and frequency of tropical cyclones in all basins to the mid-21st century. This low confidence reflects the small number of studies exploring near-term tropical cyclone activity, the differences across published projections of tropical cyclone activity, and the large role for natural variability. There is low confidence in near-term projections for increased tropical cyclone intensity in the Atlantic; this projection is in part due to projected reductions in aerosol loading. {11.3.2}
TS.5.4.5 Projected Near-Term Changes in the Ocean
It is very likely that globally-averaged surface and vertically-averaged ocean temperatures will increase in the nea-term. In the absence of multiple major volcanic eruptions, it is very likely that globally-averaged surface and depth-averaged temperatures averaged for 2016–2035 will be warmer than those averaged over 1986–2005. {11.3.3}
It is likely that salinity will increase in the tropical and (especially) subtropical Atlantic, and decrease in the western tropical Pacific over the next few decades. Overall, it is likely that there will be some decline in the Atlantic Meridional Overturning Circulation by 2050 (medium confidence). However, the rate and magnitude of weakening is very uncertain and decades when this circulation increases are also to be expected. {11.3.3}
TS.5.4.6 Projected Near-Term Changes in the Cryosphere
A nearly ice-free Arctic Ocean (sea ice extent less than 106 km2) in September is likely before mid-century under RCP8.5 (medium confidence). This assessment is based on a subset of models that most closely reproduce the climatological mean state and 1979 to 2012 trend of Arctic sea ice cover. It is very likely that there will be further shrinking and thinning of Arctic sea ice cover, and decreases of northern high-latitude spring time snow cover and near surface permafrost as global mean surface temperature rises (see Figures TS.17 and TS.18). There is low confidence in projected near-term decreases in the Antarctic sea ice extent and volume. {11.3.4}
TS.5.4.7 Possibility of Near-Term Abrupt Changes in Climate
There are various mechanisms that could lead to changes in global or regional climate that are abrupt by comparison with rates experienced in recent decades. The likelihood of such changes is generally lower for the near term than for the long term. For this reason the relevant mechanisms are primarily assessed in the TS.5 sections on long-term changes and in TFE.5. {11.3.4}
TS.5.4.8 Projected Near-Term Changes in Air Quality
There is high confidence that baseline surface ozone (O3), upon which local pollution builds, will decrease over most regions as rising temperatures enhance global O3 destruction, but it will increase with rising methane (high confidence). Projections based on the RCP scenarios differ regionally, in general projecting 2030 pollution to decrease over Europe, North America, and South America, and to increase over South Asia (medium confidence), with mixed results elsewhere. For O3, continental-scale changes across the RCPs range from –4 to +5 ppb by 2030 and –14 to +5 ppb by 2100. This range is driven more by pollutant emissions than physical climate changes (medium confidence). There is low confidence in projecting extreme pollution episodes involving ozone and particulate matter, due to difficulty projecting associated stagnation episodes that involve meteorological blocking. Nevertheless, during extreme pollution episodes, warmer temperatures can trigger positive feedbacks in chemistry and local emissions, further enhancing pollution levels (medium confidence). Future air pollution levels under the RCPs, for both surface ozone and particulate matter, are lower than under SRES scenarios with comparable greenhouse gases (high confidence). {11.3.5; Annex II}
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The source document for this Digest states:
TS.5.5 Long-Term Climate Change
TS.5.5.1 Projected Long-Term Changes in Global Temperature Global mean temperatures will continue to rise over the 21st century under all of the RCPs. From around the mid-21st century, the rate of global warming begins to be more strongly dependent on the scenario (Figure TS.15). {12.4.1}
Under the assumptions of the concentration-driven RCPs, global-mean surface temperatures for 2081–2100, relative to 1986–2005 will likely be in the 5–95% range of the CMIP5 models; 0.3°C to 1.7°C (RCP2.6), 1.1 to 2.6°C (RCP4.5), 1.4°C to 3.1°C (RCP6.0), 2.6°C to 4.8°C (RCP8.5) (see Table TS.1). With high confidence, the 5–95% range of CMIP5 is assessed as likely rather than very likely based on the assessment of TCR (see TFE.6). The 5–95% range of CMIP5 for global mean temperature change is also assessed as likely for mid-21st century, but only with medium confidence. With respect to preindustrial conditions, global temperatures averaged in the period 2081–2100 are projected to likely exceed 1.5°C above preindustrial for RCP4.5, RCP6.0 and RCP8.5 (high confidence) and are likely to exceed 2°C above preindustrial for RCP6.0 and RCP8.5 (high confidence). Temperature change above 2°C relative to preindustiral under RCP2.6 is unlikely (medium confidence). Warming above 4°C by 2081–2100 is unlikely in all RCPs (high confidence) except for RCP8.5 where it is as likely as not (medium confidence). {12.4.1; Tables 12.2, 12.3}
TS.5.5.2 Projected Long-Term Changes in Regional Temperature There is very high confidence that globally averaged changes over land will exceed changes over the ocean at the end of the 21st century by a factor that is likely in the range 1.4 to 1.7. In the absence of a strong reduction in the Atlantic Meridional Overturning, the Arctic region is projected to warm most (very high confidence) (Figure TS.15). As global mean surface temperature rises, the pattern of atmospheric zonal- mean temperatures show warming throughout the troposphere and cooling in the stratosphere, consistent with previous assessments. The consistency is especially clear in the tropical upper troposphere and the northern high latitudes {12.4.3; Box 5.1}.
It is virtually certain that, in most places, there will be more hot and fewer cold temperature extremes as global mean temperatures increase. These changes are expected for events defined as extremes on both daily and seasonal time scales. Increases in the frequency, duration and magnitude of hot extremes along with heat stress are expected, however occasional cold winter extremes will continue to occur. 20-year return values of low temperature events are projected to increase at a rate greater than winter mean temperatures in most regions, with the largest changes in the return values of low temperatures at high latitudes. 20-year return values for high temperature events are projected to increase at a rate similar to or greater than the rate of increase of summer mean temperatures in most regions. Under RCP8.5 it is likely that, in most land regions, a current 20-year high temperature event will occur more frequently by the end of the 21st century (at least doubling its frequency, but in many regions becoming an annual or two-year event) and a current 20-year low temperature event will become exceedingly rare (See also TFE.9). {12.4.3}
Models simulate a decrease in cloud amount in the future over most of the tropics and mid-latitudes, due mostly to reductions in low clouds. Changes in marine boundary layer clouds are most uncertain. Increases in cloud fraction and cloud optical depth and therefore cloud reflection are simulated in high latitudes, poleward of 50°. {12.4.3}
TS.5.5.3 Projected Long-Term Changes in Atmospheric Circulation
Mean sea level pressure is projected to decrease in high latitudes and increase in the mid-latitudes as global temperatures rise. In the tropics, the Hadley and Walker circulations are likely to slow down. Poleward shifts in the mid-latitude jets of about 1 to 2 degrees latitude are likely at the end of the 21st century under RCP8.5 in both hemispheres (medium confidence), with weaker shifts in the Northern Hemisphere. In austral summer, the additional influence of stratospheric ozone recovery in the Southern Hemisphere opposes changes due to greenhouse gases there, though the net response varies strongly across models and scenarios. Substantial uncertainty and thus low confidence remains in projecting changes in Northern Hemisphere storm tracks, especially for the North Atlantic basin. The Hadley cell is likely to widen, which translates to broader tropical regions and a poleward encroachment of subtropical dry zones. In the stratosphere, the Brewer-Dobson winds are likely to strengthen. {12.4.4}
TS.5.5.4 Projected Long-Term Changes in the Water Cycle
On the planetary scale, relative humidity is projected to remain roughly constant, but specific humidity to increase in a warming climate. The projected differential warming of land and ocean promotes changes in atmospheric moistening that lead to small decreases in near-surface relative humidity over most land areas with the notable exception of parts of tropical Africa (medium confidence) (see TFE.1, Figure 1). {12.4.5}
It is virtually certain that, in the long term, global precipitation will increase with increased global mean surface temperature. Global mean precipitation will increase at a rate per °C smaller than that of atmospheric water vapour. It will likely increase by 1 to 3% °C–1 for scenarios other than RCP2.6. For RCP2.6 the range of sensitivities in the CMIP5 models is 0.5 to 4% °C–1 at the end of the 21st century. {7.6.2, 7.6.3, 12.4.1}
Changes in average precipitation in a warmer world will exhibit substantial spatial variation under RCP8.5. Some regions will experience increases, other regions will experience decreases, and yet others will not experience significant changes at all (see Figure TS.16). There is high confidence that the contrast of annual mean precipitation between dry and wet regions and that the contrast between wet and dry seasons will increase over most of the globe as temperatures increase. The general pattern of change indicates that high latitudes are very likely to experience greater amounts of precipitation due to the increased specific humidity of the warmer troposphere as well as increased transport of water vapour from the tropics by the end of this century under the RCP8.5 scenario. Many mid-latitude and subtropical arid and semi-arid regions will likely experience less precipitation and many moist mid-latitude regions will likely experience more precipitation by the end of this century under the RCP8.5 scenario. Maps of precipitation change for the four RCP scenarios are shown in Figure TS.16. {12.4.2, 12.4.5}
Globally, for short-duration precipitation events, a shift to more intense individual storms and fewer weak storms is likely as temperatures increase. Over most of the mid-latitude land-masses and over wet tropical regions, extreme precipitation events will very likely be more intense and more frequent in a warmer world. The global average sensitivity of the 20-year return value of the annual maximum daily precipitation ranges from 4% per °C of local temperature increase (average of CMIP3 models) to 5.3% per °C of local temperature increase (average of CMIP5 models) but regionally there are wide variations. {12.4.2, 12.4.5}
Annual surface evaporation is projected to increase as global temperatures rise over most of the ocean and is projected to change over land following a similar pattern as precipitation. Decreases in annual runoff are likely in parts of southern Europe, the Middle East, and southern Africa by the end of this century under the RCP8.5 scenario. Increases in annual runoff are likely in the high northern latitudes corresponding to large increases in winter and spring precipitation by the end of this century under the RCP8.5 scenario. Regional to global-scale projected decreases in soil moisture and increased risk of agricultural drought are likely in presently dry regions and are projected with medium confidence by the end of this century under the RCP8.5 scenario. Prominent areas of projected decreases in evaporation include southern Africa and north western Africa along the Mediterranean. Soil moisture drying in the Mediterranean and southern African regions is consistent with projected changes in Hadley circulation and increased surface temperatures, so surface drying in these regions as global temperatures increase is likely with high confidence by the end of this century under the RCP8.5 scenario. In regions where surface moistening is projected, changes are generally smaller than natural variability on the twenty-year time scale. A summary of the projected changes in the water cycle from the CMIP5 models is shown in TFE.1, Figure 1. {12.4.5; Box 12.1}
TS.5.5.5 Projected Long-Term Changes in the Cryosphere
It is very likely that the Arctic sea ice cover will continue shrinking and thinning year-round in the course of the 21st century as global mean surface temperature rises. At the same time, in the Antarctic, a decrease in sea ice extent and volume is expected, but with low confidence. The CMIP5 multi-model projections give average reductions in Arctic sea ice extent for 2081‒2100 compared to 1986‒2005 ranging from 8% for RCP2.6 to 34% for RCP8.5 in February and from 43% for RCP2.6 to 94% for RCP8.5 in September (medium confidence) (Figure TS.17). A nearly ice-free Arctic Ocean (sea ice extent less than 106 km2) in September before mid-century is likely under RCP8.5 (medium confidence), based on an assessment of a subset of models that most closely reproduce the climatological mean state and 1979‒2012 trend of the Arctic sea ice cover. Some climate projections exhibit 5 to 10 year periods of sharp summer Arctic sea ice decline ‒ even steeper than observed over the last decade ‒ and it is likely that such instances of rapid ice loss will occur in the future. There is little evidence in global climate models of a tipping point (or critical threshold) in the transition from a perennially ice-covered to a seasonally ice-free Arctic Ocean beyond which further sea ice loss is unstoppable and irreversible. In the Antarctic, the CMIP5 multi-model mean projects a decrease in sea ice extent that ranges from 16% for RCP2.6 to 67% for RCP8.5 in February and from 8% for RCP2.6 to 30% for RCP8.5 in September for 2081‒2100 compared to 1986‒2005. There is however low confidence in those projections because of the wide inter-model spread and the inability of almost all of the available models to reproduce the overall increase of the Antarctic sea ice areal coverage observed during the satellite era. {12.4.6, 12.5.5}
It is very likely that Northern Hemisphere snow cover will reduce as global temperatures rise over the coming century. A retreat of permafrost extent with rising global temperatures is virtually certain. Snow cover changes result from precipitation and ablation changes, which are sometimes opposite. Projections of the Northern Hemisphere spring snow covered area by the end of the 21st century vary between –7% (RCP2.6) and –25% (RCP8.5) decrease and are fairly coherent among models (Figure TS.18). The projected changes in permafrost are a response not only to warming, but also to changes in snow cover, which exerts a control on the underlying soil. By the end of the 21st century, diagnosed near-surface permafrost area is projected to decrease by between 37% (RCP2.6) to 81% (RCP8.5) (medium confidence). {12.4.6}
TS.5.5.6 Projected Long-Term Changes in the Ocean
Over the course of the 21st century, the global ocean will warm in all RCP scenarios. The strongest warming signal is found at the surface in subtropical and tropical regions. At greater depth the warming is most pronounced in the Southern Ocean. In some regions, ocean warming in the top few hundred meters can exceed 0.5°C (RCP2.6) to 2.5°C (RCP8.5), and 0.3°C (RCP2.6) to 0.7°C (RCP8.5) at a depth of about 1 km by the end of the century. For RCP4.5 by the end of the 21st century, half of the energy taken up by the ocean is in the uppermost 700 m, and 85% is in the uppermost 2000 m. Due to the long time scales of this heat transfer from the surface to depth, ocean warming will continue for centuries, even if greenhouse gas emissions are decreased or concentrations kept constant, and will result in a continued contribution to sea level rise (see Section TS5.7). {12.4.3, 12.4.7}
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TS.5.7 Long-Term Projections of Sea Level Change
TS.5.7.1 Projections of Global Mean Sea Level Change for the 21st Century Global mean sea level rise for 2081–2100 (relative to 1986–2005) for the RCPs will likely be in the 5–95% ranges derived from CMIP5 climate projections in combination with process-based models of glacier and ice-sheet SMB, with possible ice-sheet dynamical changes assessed from the published literature, These likely ranges are 0.26 to 0.54 m (RCP2.6), 0.32 to 0.62 m (RCP4.5), 0.33 to 0.62 m (RCP6.0), 0.45 to 0.81 (RCP8.5) m (medium confidence) (see Table TS.1 and Figure TS.21). For RCP8.5 the range at 2100 is 0.53 to 0.97 m. The central projections for global mean sea level rise in all scenarios lie within a range of 0.05 m until the middle of the century, when they begin to diverge; by the end of the century, they have a spread of about 0.3 m. Although RCP4.5 and RCP6.0 are very similar at the end of the century, RCP4.5 has a greater rate of rise earlier in the century than RCP6.0. GMSL rise depends on the pathway of CO2 emissions, not only on the cumulative total; reducing emissions earlier rather than later, for the same cumulative total, leads to a larger mitigation of sea level rise. {12.4.1, 13.4.1, 13.5.1; Table 13.5}
Confidence in the projected likely ranges comes from the consistency of process-based models with observations and physical understanding. The basis for higher projections has been considered and it has been concluded that there is currently insufficient evidence to evaluate the probability of specific levels above the likely range. Based on current understanding, only the collapse of marine-based sectors of the Antarctic ice sheet, if initiated, could cause GMSL to rise substantially above the likely range during the 21st century. There is a lack of consensus on the probability for such a collapse, and the potential additional contribution to GMSL rise cannot be precisely quantified, but there is medium confidence that it would not exceed several tenths of a meter of sea level rise during the 21st century. {13.5.1, 13.5.3}
Under all the RCP scenarios, the time-mean rate of global mean sea level rise during the 21st century is very likely to exceed the rate observed during 1971–2010. In the projections, the rate of rise initially increases. In RCP2.6 it becomes roughly constant (central projection ~4.5 mm yr–1) before the middle of the century, and subsequently declines slightly. The rate of rise becomes roughly constant in RCP4.5 and RCP6.0 by the end of the century, whereas acceleration continues throughout the century in RCP8.5 (reaching 11 [7 to 15] mm yr–1 during 2081–2100). {13.5.1; Table 13.5}
In all RCP scenarios, thermal expansion is the largest contribution, accounting for about 30–55% of the total. Glaciers are the next largest. By 2100, 15–55% of the present glacier volume is projected to be eliminated under RCP2.6, and 35 to 85% under RCP8.5 (medium confidence). The increase in surface melting in Greenland is projected to exceed the increase in accumulation, and there is high confidence that the surface mass balance changes on the Greenland ice sheet will make a positive contribution to sea level rise over the 21st century. On the Antarctic ice sheet, surface melting is projected to remain small, while there is medium confidence that snowfall will increase (Figure TS.21). {13.3.3, 13.4.3, 13.4.4, 13.5.1; Table 13.5}
There is medium confidence in the ability to model future rapid changes in ice-sheet dynamics on decadal timescales. At the time of the AR4, scientific understanding was not sufficient to allow an assessment of the possibility of such changes. Since the publication of the AR4, there has been substantial progress in understanding the relevant processes as well as in developing new ice-sheet models that are capable of simulating them. However, the published literature as yet provides only a partially sufficient basis for making projections related to particular scenarios. In our projections of GMSL rise by 2081–2100, the likely range from rapid changes in ice outflow is 0.03 to 0.20 m from the two ice sheets combined, and its inclusion is the most important reason why the projections are greater than those given in the AR4. {13.1.5, 13.5.1, 13.5.3}
Semi-empirical models are designed to reproduce the observed sea level record over their period of calibration, but do not attribute sea level rise to its individual physical components. For RCPs, some semi- empirical models project a range that overlaps the process-based likely range while others project a median and 95-percentile that are about twice as large as the process-based models. In nearly every case, the semi- empirical model 95-percentile is higher than the process-based likely range. For 2081–2100 (relative to 1986–2005) under RCP4.5, semi-empirical models give median projections in the range 0.56 to 0.97 m, and their 95-percentiles extend to about 1.2 m. This difference implies either that there is some contribution which is presently unidentified or underestimated by process-based models, or that the projections of semi- empirical models are overestimates. Making projections with a semi-empirical model assumes that sea level change in the future will have the same relationship as it has had in the past to radiative forcing or global mean temperature change. This may not hold if potentially non-linear physical processes do not scale in the future in ways which can be calibrated from the past. There is no consensus in the scientific community about the reliability of SEM projections, and confidence in them is assessed to be low. {13.5.2, 13.5.3}
TS.5.7.2 Projections of Global Mean Sea Level Change Beyond 2100
It is virtually certain that global mean sea level rise will continue beyond 2100. The few available model results that go beyond 2100 indicate global mean sea level rise above the pre-industrial level by 2300 to be less than 1 m for a radiative forcing that corresponds to CO2 concentrations that peak and decline and remain below 500 ppm, as in the scenario RCP2.6. For a radiative forcing that corresponds to a CO2 concentration that is above 700 ppm but below 1500 ppm, as in the scenario RCP8.5, the projected rise is 1 m to more than 3 m (medium confidence). {13.5} Sea level rise due to ocean thermal expansion will continue for centuries to millennia. The amount of ocean thermal expansion increases with global warming (models give a range of 0.2 to 0.6 m °C–1). The glacier contribution decreases over time as their volume (currently ~0.43 m sea level equivalent) decreases. In Antarctica, beyond 2100 and with higher greenhouse gas scenarios, the increase in surface melting could exceed the increase in accumulation. {13.5.2, 13.5.4}
The available evidence indicates that global warming greater than a certain threshold would lead to the near- complete loss of the Greenland Ice Sheet over a millennium or more, causing a global mean sea level rise of about 7 m. Studies with fixed present-day ice sheet topography indicate the threshold is greater than 2°C but less than 4°C of global mean surface temperature rise with respect to preindustrial. Taking into account the increased vulnerability of the ice sheet as the surface elevation decreases due to the loss of ice, a study with a dynamical ice sheet suggests the threshold could be as low as 1°C. Considering the present state of scientific uncertainty, a likely range cannot be quantified. The complete loss of the ice sheet is not inevitable because this would take a millennium or more; if temperatures decline before the ice sheet is eliminated, the ice sheet might regrow. However, some part of the mass loss might be irreversible, depending on the duration and degree of exceedance of the threshold, because the ice sheet may have multiple steady states, due to its interaction with its regional climate. {13.4.3, 13.5.4}
Currently available information indicates that the dynamical contribution of the ice sheets will continue beyond 2100, but confidence in projections is low. In Greenland, ice outflow induced from interaction with the ocean is self-limiting as the ice-sheet margin retreats inland from the coast. By contrast, the bedrock topography of Antarctica is such that there may be enhanced rates of mass loss as the ice retreats. About 3.3 m of equivalent global sea level of the West Antarctic ice sheet is grounded on areas with downward sloping bedrock, which may be subject to potential ice loss via the marine ice-sheet instability. Due to relatively weak snowfall on Antarctica and the slow ice motion in its interior, it can be expected that the West Antarctic ice sheet would take at least several thousand years to regrow if it was eliminated by dynamic ice discharge. Consequently any significant ice loss from West Antarctic that occurs within the next century will be irreversible on a multi-centennial to millennial time scale. {13.4.4, 13.5.4}
TS.5.7.3 Projections of Regional Sea Level Change
Regional sea level will change due to dynamical ocean circulation changes, changes in the heat content of the ocean, mass redistribution in the entire Earth system, and changes in atmospheric pressure. Ocean dynamical change results from changes in wind and buoyancy forcing (heat and freshwater), associated changes in the circulation, and redistribution of heat and freshwater. Over timescales longer than a few days, regional sea level also adjusts nearly isostatically to regional changes in sea level atmospheric pressure relative to its mean over the ocean. Ice-sheet mass loss (both contemporary and past), glacier mass loss and changes in terrestrial hydrology cause water mass redistribution among the cryosphere, the land and the oceans, giving rise to distinctive regional changes in the solid Earth, Earth rotation and the gravity field. In some coastal locations, changes in the hydrological cycle, ground subsidence associated with anthropogenic activity, tectonic processes, and coastal processes can dominate the relative sea level change, i.e., the change in sea surface height relative to the land. {13.1.3, 13.6.2, 13.6.3, 13.6.4}
By the end of the 21st century, sea level change will have a strong regional pattern, which will dominate over variability, with many regions likely experiencing substantial deviations from the global mean change (Figure TS.23). It is very likely that over about 95% of the ocean will experience regional sea level rise, while most regions experiencing a sea level fall are located near current and former glaciers and ice sheets. Local sea level changes deviate more than 10% and 25% from the global mean projection for as much as 30% and 9% of the ocean area, respectively, indicating that spatial variations can be very large. Regional changes in sea level reach values of up to 30% above the global mean value in the Southern Ocean and around North America, between 10% to 20% in equatorial regions, and up to 50% below the global mean in the Arctic region and some regions near Antarctica. About 70% of the coastlines worldwide are projected to experience a sea level change within 20% of the global mean sea level change. Over decadal periods, the rates of regional sea level change as a result of climate variability can differ from the global average rate by more than 100%. {13.6.5}
TS.5.7.4 Projections of Change in Sea Level Extremes and Waves During the 21st Century
It is very likely that there will be a significant increase in the occurrence of future sea level extremes (see also TFE.9, Table 1). This increase will primarily be the result of an increase in mean sea level (high confidence), with extreme return periods decreasing by at least an order of magnitude in some regions by the end of the 21st century. There is low confidence in region-specific projections of storminess and associated storm surges. {13.7.2}
It is likely (medium confidence) that annual mean significant wave heights will increase in the Southern Ocean as a result of enhanced wind speeds. Southern-Ocean generated swells are likely to affect heights, periods, and directions of waves in adjacent basins. It is very likely that wave heights and the duration of the wave season will increase in the Arctic Ocean as a result of reduced sea-ice extent. In general, there is low confidence in region-specific projections due to the low confidence in tropical and extratropical storm projections, and to the challenge of down-scaling future wind states from coarse resolution climate models. {13.7.3}
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TS.5.8 Climate Phenomena and Regional Climate Change
This section assesses projected changes over the 21st century in large-scale climate phenomena that affect regional climate (Table TS.2). Some of these phenomena are defined by climatology (e.g., monsoons), and some by interannual variability (e.g., El Niño), the latter affecting climate extremes such as floods, droughts and heat waves. Changes in statistics of weather phenomena such as tropical cyclones and extratropical storms are also summarized here. {14.8}
Table TS.2: Overview of projected regional changes and their relation to major climate phenomena.
TS.5.8.1 Monsoon Systems
Global measures of monsoon by the area and summer precipitation are likely to increase in the 21st century, while the monsoon circulation weakens. Monsoon onset dates are likely to become earlier or not to change much while monsoon withdrawal dates are very likely to delay, resulting in a lengthening of the monsoon season (Figure TS.24). The increase in seasonal-mean precipitation is pronounced in the East and South Asian summer monsoons while the change in other monsoon regions is subject to larger uncertainties. {14.2.1}
There is medium confidence that monsoon-related interannual rainfall variability will increase in the future. Future increase in precipitation extremes related to the monsoon is very likely in South America, Africa, East Asia, South Asia, Southeast Asia and Australia. {14.2.1, 14.8.5, 14.8.7, 14.8.9, 14.8.11, 14.8.12, 14.8.13}
There is medium confidence that overall precipitation associated with the Asian-Australian monsoon will increase but with a north-south asymmetry: Indian monsoon rainfall is projected to increase, while projected changes in the Australian summer monsoon rainfall are small. There is medium confidence in that the Indian summer monsoon circulation weakens, but this is compensated by increased atmospheric moisture content, leading to more rainfall. For the East Asian summer monsoon, both monsoon circulation and rainfall are projected to increase. {14.2.2, 14.8.9, 14.8.11, 14.8.13}
There is low confidence in projections of the North American and South American monsoon precipitation changes, but medium confidence that the North American monsoon will arrive and persist later in the annual cycle, and high confidence in expansion of South American Monsoon area. {14.2.3, 14.8.3, 14.8.4; 14.8.5}
There is low confidence in projections of a small delay in the West African rainy season, with an intensification of late-season rains. The limited skills of model simulations for the region suggest low confidence in the projections. {14.2.4, 14.8.7}
TS.5.8.2 Tropical Phenomena
It is virtually certain that precipitation change will vary in space, increasing in some regions and decreasing in some others. The spatial distribution of tropical rainfall changes is likely shaped by the current climatology and ocean warming pattern. The first effect is to increase rainfall near the currently rainy regions, and the second effect increases rainfall where the ocean warming exceeds the tropical mean. There is medium confidence that tropical rainfall projections are more reliable for the seasonal than annual-mean changes. {7.6.2, 12.4.5, 14.3.1}
There is medium confidence in future increase in seasonal-mean precipitation on the equatorial flank of the Inter-Tropical Convergence Zone and a decrease in precipitation in the subtropics including parts of North and Central Americas, the Caribbean, South America, Africa and West Asia. There is medium confidence that the interannual occurrence of zonally-oriented South Pacific Convergence Zone events will increase, leading possibly to more frequent droughts in the southwest Pacific. There is medium confidence that the South Atlantic Convergence Zone will shift southwards, leading to a precipitation increase over south- eastern South America and a reduction immediately north of the convergence zone. {14.3.1, 14.8.3, 14.8.4, 14.8.5, 14.8.7, 14.8.11, 14.8.14}
The tropical Indian Ocean is likely to feature a zonal pattern with reduced warming and decreased rainfall in the east (including Indonesia), and enhanced warming and increased rainfall in the west (including East Africa). The Indian Ocean dipole mode of interannual variability is very likely to remain active, affecting climate extremes in East Africa, Indonesia, and Australia. {14.3.3, 14.8.7, 14.8.12}
There is low confidence in the projections for the tropical Atlantic - both for the mean and interannual modes, because of large errors in model simulations in the region. Future projections in Atlantic hurricanes, tropical South American and West African precipitation are therefore of low confidence. {14.3.4, 14.6.1, 14.8.5 and 14.8.7}
There is low confidence in projections of future changes in the Madden-Julian Oscillation due to the poor skill in model simulations of this intraseasonal phenomenon and the sensitivity to ocean warming patterns. Future projections of regional climate extremes in West Asia, Southeast Asia and Australia are therefore of low confidence. {9.5.2, 14.3.4, 14.8.10, 14.8.12, 14.8.13}
TS.5.8.3 El Niño-Southern Oscillation
El Niño-Southern Oscillation very likely remains as a dominant mode of interannual variability in the future and regional rainfall variability it induces likely intensifies. Natural variations of the amplitude and spatial pattern of El Niño-Southern Oscillation are so large that confidence in any projected change for the 21st century remains low. The projected change in El Niño amplitude is small for both RCP4.5 and RCP8.5 compared to the spread of the change among models (Figure TS.25). Over the North Pacific and North America, patterns of temperature and precipitation anomalies related to El Niño and La Niña (teleconnections) are likely to move eastwards in the future, while confidence is low in changes in climate impacts on other regions including Central and South Americas, the Caribbean, Africa, most of Asia, Australia and most Pacific Islands. {14.4, 14.8.3, 14.8.4, 14.8.5, 14.8.7, 14.8.9, 14.8.11, 14.8.12, 14.8.13, 14.8.14}
TS.5.8.4 Cyclones
Projections for the 21st century indicate that it is likely that the global frequency of tropical cyclones will either decrease or remain essentially unchanged, concurrent with a likely increase in both global mean tropical cyclone maximum wind speed and rain rates (Figure TS.26). The influence of future climate change on tropical cyclones is likely to vary by region, but there is low confidence in region-specific projections. The frequency of the most intense storms will more likely than not increase substantially in the Western North Pacific and North Atlantic. More extreme precipitation near the centers of tropical cyclones making landfall are likely in North and Central America, East Africa, West, East, South and Southeast Asia as well as in Australia and many Pacific islands. {14.6.1, 14.8.3, 14.8.4, 14.8.7, 14.8.9, 14.8.10, 14.8.11, 14.8.12, 14.8.13, 14.8.14}
The global number of extra-tropical cyclones is unlikely to decrease by more than a few percent and future changes in storms are likely to be small compared to natural interannual variability and substantial variations between models. A small poleward shift is likely in the Southern Hemisphere storm track but the magnitude of this change is model-dependent. It is unlikely that the response of the North Atlantic storm track in climate projections is a simple poleward shift. There is medium confidence in a projected poleward shift in the North Pacific storm track. There is low confidence in the impact of storm track changes on regional climate at the surface . More precipitation in extra-tropical cyclones leads to a winter precipitation increase in Arctic, Northern Europe, North America, and the mid-high latitude Southern Hemisphere. {11.3.2, 12.4.4, 14.6.2, 14.8.2, 14.8.3, 14.8.5, 14.8.6, 14.8.13, 14.8.15}
TS.5.8.5 Annular and Dipolar Modes of Variability
Future boreal wintertime North Atlantic Oscillation is very likely to exhibit large natural variations as observed in the past. The North Atlantic Oscillation is likely to become slightly more positive (on average), with some, but not very well documented implications for winter conditions in the Arctic, North America and Eurasia. The austral summer/autumn positive trend in Southern Annular Mode is likely to weaken considerably as stratospheric ozone recovers through the mid-21st century with some, but not very well documented, implications for South America, Africa, Australia, New Zealand and Antarctica. {11.3.2, 14.5.2 and 14.8.5, 14.8.7, 14.8.13, 14.8.15}
TS.5.8.6 Additional Phenomena
It is unlikely that the Atlantic Multi-Decadal Oscillation will change its behaviour as the mean climate changes. However, natural fluctuations in the Atlantic Multi-Decadal Oscillation over the coming few decades are likely to influence regional climates at least as strongly as will human-induced changes with implications for Atlantic major hurricane frequency, the West African monsoon, North American and European summer conditions. {14.2.4, 14.5.1, 14.6.1, 14.7.6, 14.8.2, 14.8.3, 14.8.6, 14.8.8}
There is medium confidence that the frequency of Northern and Southern Hemisphere blocking will not increase, but confidence is low in the trends in blocking intensity and persistence. {Box 14.2}
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TFE.8: Climate Targets and Stabilization
The concept of stabilization is strongly linked to the ultimate objective of the UNFCCC, which is “to achieve [...] stabilization of greenhouse gas concentrations in the atmosphere at a level that would prevent dangerous anthropogenic interference with the climate system.” Recent policy discussions focused on limits to a global temperature increase, rather than to greenhouse gas concentrations, as climate targets in the context of the UNFCCC objectives. The most widely discussed is that of 2°C, i.e., to limit global temperature increase relative to preindustrial times to below 2°C, but targets other than 2°C have been proposed (e.g., returning warming to well below 1.5°C global warming relative to preindustrial, or returning below an atmospheric CO2 concentration of 350 ppm). Climate targets generally mean avoiding a warming beyond a predefined threshold. Climate impacts however are geographically diverse and sector specific, and no objective threshold defines when dangerous interference is reached. Some changes may be delayed or irreversible, and some impacts could be beneficial. It is thus not possible to define a single critical objective threshold without value judgments and without assumptions on how to aggregate current and future costs and benefits. This section does not advocate or defend any threshold or objective, nor does it judge the economic or political feasibility of such goals, but assesses, based on the current understanding of climate and carbon cycle feedbacks, the climate projections following the RCP scenarios in the context of climate targets, and the implications of different long-term temperature stabilization objectives on allowed carbon emissions. Further below it is highlighted that temperature stabilization does not necessarily imply stabilization of the entire Earth system.
Temperature targets imply an upper limit on the total radiative forcing (RF). Differences in RF between the four RCP scenarios are relatively small up to 2030, but become very large by the end of the 21st century and dominated by CO2 forcing. Consequently, in the near term, global-mean surface temperatures are projected to continue to rise at a similar rate for the four RCP scenarios. Around the mid-21st century, the rate of global warming begins to be more strongly dependent on the scenario. By the end of the 21st century, global mean temperatures will be warmer than present day under all the RCPs, global temperature change being largest (>0.3°C per decade) in the highest RCP8.5 and significantly lower in RCP2.6, particularly after ~2050 when global surface temperature response stabilizes (and declines thereafter) (see Figure TS.15). {11.3.1, 12.3.3, 12.4.1}
In the near term (2016-2035), global mean surface warming is more likely than not to exceed 1°C and very unlikely to be more than 1.5oC relative to preindustrial (assuming 0.61°C warming has occurred prior to 1986–2005) (medium confidence). By the end of the century (2081–2100), global mean surface warming, relative to preindustrial, is likely to exceed 1.5°C for RCP4.5, RCP6.0 and RCP8.5 (high confidence) and is likely to exceed 2°C for RCP6.0 and RCP8.5 (high confidence). Global mean surface warming above 2°C under RCP2.6 is unlikely (medium confidence). Global mean surface warming above 4°C by 2081–2100 is unlikely in all RCPs (high confidence) except for RCP8.5 where it is as likely as not (medium confidence). All targets here are defined as global average surface temperature change relative to preindustrial. {11.3.6, 12.4.1; Table 12.3}
Continuing greenhouse gas emissions beyond 2100 as in the RCP8.5 extension induces a total radiative forcing above 12 W m–2 by 2300, global warming reaching 7.8°C [3.0 to 12.6°C] for 2281–2300 relative to 1986–2005. Under the RCP4.5 extension, where radiative forcing is kept constant (around 4.5 W m-2) beyond 2100, global warming reaches 2.5°C [1.5 to 3.5°C]. Global warming reaches 0.6°C [0.0 to 1.2°C] under the RCP2.6 extension where sustained negative emissions lead to a further decrease in radiative forcing, reaching values below present-day radiative forcing by 2300. See also Box TS.7. {12.3.1, 12.4.1, 12.5.1}
The total amount of anthropogenic CO2 released in the atmosphere since preindustrial (often termed cumulative carbon emission, although it only applies to CO2 emissions) is a good indicator of the atmospheric CO2 concentration and hence of the global warming response. The ratio of global mean surface temperature change to total cumulative anthropogenic CO2 emissions is relatively constant over time and independent of the scenario. This near-linear relationship between total CO2 emissions and global temperature change makes it possible to define a new quantity, the transient climate response to cumulative carbon emission (TCRE), as the transient global mean surface temperature change for a given amount of cumulated anthropogenic CO2 emissions, usually 1000 PgC. TCRE is model dependent, as it is a function of the cumulative CO2 airborne fraction and the transient climate response, both quantities varying significantly across models. Taking into account the available information from multiple lines of evidence (observations, models and process understanding), the near linear relationship between cumulative CO2 emissions and peak global mean temperature is well established in the literature and robust for cumulative total CO2 emissions up to about 2000 PgC. It is consistent with the relationship inferred from past cumulative CO2 emissions and observed warming, is supported by process understanding of the carbon cycle and global energy balance, and emerges as a robust result from the entire hierarchy of models. Expert judgment based on the available evidence suggests that TCRE is likely between 0.8°C–2.5°C per 1000 PgC, for cumulative emissions less than about 2000 PgC until the time at which temperature peaks (TFE.8, Figure 1a). {6.4.3, 12.5.4; Box 12.2}
CO2 induced warming is projected to remain approximately constant for many centuries following a complete cessation of emissions. A large fraction of climate change is thus irreversible on a human time scale, except if net anthropogenic CO2 emissions were strongly negative over a sustained period. The assessment of TCRE, limiting the warming caused by anthropogenic CO2 emissions alone to be likely less than 2°C, total CO2 emissions from all anthropogenic sources would need to be below a cumulative budget of about 1000 PgC over the entire industrial era. About half, estimated in the range of 460 to 630 PgC, of this budget was already emitted by 2011 (TFE.8, Figure 1a). Higher emissions in earlier decades therefore imply lower or even negative emissions later on. Accounting for non-CO2 forcings contributing to peak warming implies lower cumulated CO2 emissions (TFE.8, Figure 1a). Non-CO2 forcing constituents are important, requiring either assumptions on how CO2 emission reductions are linked to changes in other forcings, or separate emission budgets and climate modeling for short lived and long-lived gases. So far, not many studies have considered non-CO2 forcings. Those that do consider them found significant effects, in particular warming of several tenths of a degree for abrupt reductions in emissions of short-lived species, like aerosols. Accounting for an unanticipated release of greenhouse gases from permafrost or methane hydrates, not included in studies assessed here, would also reduce the anthropogenic CO2 emissions compatible with a given temperature target. Requiring a higher likelihood of temperatures remaining below a given temperature target would further reduce the compatible emissions (TFE.8, Figure 1c). The concept of a fixed cumulative CO2 budget holds not just for 2°C, but for any temperature level explored with models so far (up to about 5°C, see Figures 12.44–12.46). Higher temperature targets would allow larger cumulative budgets, while lower temperature target would require lower cumulative budgets (TFE.8, Figure 1). {6.3.1, 12.5.2, 12.5.4}
The climate system has multiple timescales, ranging from annual to multi-millennial, associated with different thermal and carbon reservoirs. These long time scales induce a commitment warming “already in the pipe-line”. Stabilization of the forcing would not lead to an instantaneous stabilization of the warming. For the RCP scenarios and their extensions to 2300, the fraction of realized warming, at that time when radiative forcing stabilizes, would be about 75 to 85% of the equilibrium warming. For a 1% yr–1 CO2 increase to 2 × CO2 or 4 × CO2 and constant forcing thereafter, the fraction of realized warming would much smaller, about 40 to 70% at the time when the forcing is kept constant. Due to the long timescales in the deep ocean, full equilibrium is reached only after hundreds to thousands of years.
The commitment to past emissions is a persistent warming for hundreds of years, continuing at about the level of warming that has been realized when emissions were ceased. The persistence of this CO2-induced warming after emission have ceased results from a compensation between the delayed commitment warming described above and the slow reduction in atmospheric CO2 resulting from ocean and land carbon uptake. This persistence of warming also results from the nonlinear dependence of radiative forcing on atmospheric CO2, i.e., the relative decrease in forcing being smaller than the relative decrease in CO2 concentration. For high climate sensitivities, and in particular if sulphate aerosol emissions are eliminated at the same time as greenhouse gas emissions, the commitment from past emission can be strongly positive, and is a superposition of a fast response to reduced aerosols emissions and a slow response to reduced CO2. {12.5.4}
Stabilization of global temperature does not imply stabilization for all aspects of the climate system. Processes related to vegetation change, changes in the ice sheets, deep ocean warming and associated sea level rise and potential feedbacks linking for example ocean and the ice sheets have their own intrinsic long timescales. Ocean acidification will very likely continue in the future as long as the oceans will continue to take up atmospheric CO2. Committed land ecosystem carbon cycle changes will manifest themselves further beyond the end of the 21st century. It is virtually certain that global mean sea level rise will continue beyond 2100, with sea level rise due to thermal expansion to continue for centuries to millennia. Global mean sea level rise depends on the pathway of CO2 emissions, not only on the cumulative total; reducing emissions earlier rather than later, for the same cumulative total, leads to a larger mitigation of sea level rise. {6.4.4, 12.5.4, 13.5.4}
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Box TS.7: Climate Geoengineering Methods
Geoengineering is defined as the deliberate large-scale intervention in the Earth system to counter undesirable impacts of climate change on the planet. Carbon Dioxide Reduction (CDR) aims to slow or perhaps reverse projected increases in the future atmospheric CO2 concentrations, accelerating the natural removal of atmospheric CO2 and increasing the storage of carbon in land, ocean and geological reservoirs. Solar Radiation Management (SRM) aims to counter the warming associated with increasing greenhouse gas concentrations by reducing the amount of sunlight absorbed by the climate system. A related technique seeks to deliberately decrease the greenhouse effect in the climate system by altering high-level cloudiness. {6.5, 7.7; FAQ 7.3}
CDR methods could provide mitigation of climate change if CO2 can be reduced, but there are uncertainties, side effects and risks, and implementation would depend on technological maturity along with economic, political and ethical considerations. CDR would need to be deployed at large-scale and over at least one century to be able to significantly reduce CO2 concentrations. There are biogeochemical, and currently technical limitations that make it difficult to provide quantitative estimates of the potential for CDR. CO2 removals from the atmosphere by CDR would be partially offset by outgassing of CO2 previously stored in ocean and terrestrial carbon reservoirs. Some of the climatic and environmental side effects of CDR methods are associated with altered surface albedo from afforestation, ocean de-oxygenation from ocean fertilization, and enhanced N2O emissions. Land-based CDR methods would probably face competing demands for land. The level of confidence on the effectiveness of CDR methods and their side effects on carbon and other biogeochemical cycles is low. {6.5; Box 6.2; FAQ 7.3}
Solar Radiation Management (SRM) remains unimplemented and untested but, if realisable, could offset a global temperature rise and some of its effects. There is medium confidence that SRM through stratospheric aerosol injection is scalable to counter the radiative forcing (RF) and some of the climate effects expected from a twofold increase in CO2 concentration. There is no consensus on whether a similarly large RF could be achieved from cloud brightening SRM due to insufficient understanding of aerosol-cloud interactions. It does not appear that land albedo change SRM could produce a large RF. Limited literature on other SRM methods precludes their assessment. {7.7.2, 7.7.3}
Numerous side effects, risks and shortcomings from SRM have been identified. SRM would produce an inexact compensation for the RF by greenhouse gases. Several lines of evidence indicate that SRM would produce a small but significant decrease in global precipitation (with larger differences on regional scales) if the global surface temperature were maintained. Another side effect that is relatively well characterised is the likelihood of modest polar stratospheric ozone depletion associated with stratospheric aerosol SRM. There could also be other as yet unanticipated consequences {7.6.3, 7.7.3, 7.7}
As long as greenhouse gas concentrations continued to increase, the SRM would require commensurate increase, exacerbating side effects. Additionally, scaling SRM to substantial levels would carry the risk that if the SRM were terminated for any reason, there is high confidence that surface temperatures would increase rapidly (within a decade or two) to values consistent with the greenhouse gas forcing, which would stress systems sensitive to the rate of climate change. Finally, SRM would not compensate for ocean acidification from increasing CO2. {7.7.3, 7.7.4}
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TFE.9: Climate Extremes
Assessing changes in climate extremes poses unique challenges, not just because of the intrinsically rare nature of these events, but because they invariably happen in conjunction with disruptive conditions. They are strongly influenced by both small and large-scale weather patterns, modes of variability, thermodynamic processes, land-atmosphere feedbacks and antecedent conditions. Much progress has been made since AR4 including the comprehensive assessment of extremes undertaken by the IPCC Special Report on Managing the Risk of Extreme Events and Disasters to Advance Climate Change Adaptation (SREX) but also due to the amount of observational evidence available, improvements in our understanding and the ability of models to simulate extremes. {1.3.3, 2.6, 7.6, 9.5.4}
For some climate extremes such as droughts, floods and heat waves, several factors need to be combined to produce an extreme event. Analyses of rarer extremes such as 1-in-20 to 1-in-100 year events using Extreme Value Theory are making their way into a growing body of literature. Other recent advances concern the notion of “fraction of attributable risk” that aims to link a particular extreme event to specific causal relationships. {1.3.3, 2.6.1, 2.6.2, 10.6.2, 12.4.3; Box 2.4}
TFE.9, Table 1 indicates the changes that have been observed in a range of weather and climate extremes over the last 50 years, the assessment of the human contribution to those changes, and how those extremes are expected to change in the future. The table also compares the current assessment with that of the AR4 and the SREX where applicable. {2.6, 3.7, 10.6, 11.3, 12.4, 14.6}
Temperature Extremes, Heat Waves and Warm Spells
It is very likely that both maximum and minimum temperature extremes have warmed over most land areas since the mid-20th century. These changes are well simulated by current climate models, and it is very likely that anthropogenic forcing has affected the frequency of these extremes and virtually certain that further changes will occur. This supports AR4 and SREX conclusions although with greater confidence in the anthropogenic forcing component. {2.6.1, 9.5.4, 10.6.1, 12.4.3}
For land areas with sufficient data there has been an overall increase in the number of warm days and nights. Similar decreases are seen in the number of cold days and nights. It is very likely that increases in unusually warm days and nights and/or reductions in unusually cold days and nights including frosts have occurred over this period across most continents. Warm spells or heat waves containing consecutive extremely hot days or nights are often associated with quasi-stationary anticyclonic circulation anomalies and are also affected by pre-existing soil conditions and the persistence of soil moisture anomalies that can amplify or dampen heat waves particularly in moisture-limited regions. Most global land areas, with a few exceptions, have experienced more heat waves since the middle of the 20th century. Several studies suggest that increases in mean temperature account for most of the changes in heat wave frequency, however, heat wave intensity/amplitude is highly sensitive to changes in temperature variability and the shape of the temperature distribution and heat wave definition also plays a role. While in some regions instrumental periods prior to the 1950s had more heat waves (e.g., USA), for other regions such as Europe, an increase in heat wave frequency in the period since the 1950s stands out in long historical temperature series. {2.6, 2.6.1, 5.5.1; Box 2.4; Table 2.12, 2.13; FAQ 2.2} The observed features of temperature extremes and heat waves are well simulated by climate models and are similar to the spread amongst observationally-based estimates in most regions. Regional downscaling now offers credible information on the spatial scales required for assessing extremes and improvements in the simulation of the El Nino-Southern Oscillation from CMIP3 to CMIP5 and other large-scale phenomenon is crucial. However simulated changes in frequency and intensity of extreme events is limited by observed data availability and quality issues and by the ability of models to reliably simulate certain feedbacks and mean changes in key features of circulation such as blocking. {2.6, 2.7, 9.4, 9.5.3, 9.5.4, 9.6, 9.6.1, 10.3, 10.6, 14.4; Box 14.2}
Since AR4, the understanding of mechanisms and feedbacks leading to changes in extremes has improved. There continues to be strengthening evidence for a human influence on the observed frequency of extreme temperatures and heat waves in some regions. Near-term (decadal) projections suggest likely increases in temperature extremes but with little distinguishable separation between emissions scenarios (TFE.9, Figure 1). Changes may proceed at a different rate than the mean warming however, with several studies showing that projected European high-percentile summer temperatures will warm faster than mean temperatures. Future changes associated with the warming of temperature extremes in the long-term are virtually certain and scale with the strength of emissions scenario i.e., greater anthropogenic emissions correspond to greater warming of extremes (TFE.9, Figure 1). For high emissions scenarios, it is likely that, in most land regions, a current one-in-20 year maximum temperature event will at least double in frequency but in many regions will become an annual or a one-in-two year event by the end of the 21st century. The magnitude of both high and low temperature extremes is expected to increase at least at the same rate as the mean, but with 20-year return values for low temperature events projected to increase at a rate greater than winter mean temperatures in most regions. {10.6.1, 11.3.2, 12.4.3}
Precipitation Extremes
It is likely that the number of heavy precipitation events over land has increased in more regions than it has decreased in since the mid-20th century, and there is medium confidence that anthropogenic forcing has contributed to this increase. {2.6.2, 10.6.1} There has been substantial progress between CMIP3 and CMIP5 in the ability of models to simulate more realistic precipitation extremes. However, evidence suggests that the majority of models underestimate the sensitivity of extreme precipitation to temperature variability or trends especially in the tropics which implies that models may underestimate the projected increase in extreme precipitation in the future. While progress has been made in understanding the processes that drive extreme precipitation, challenges remain in quantifying cloud and convective effects in models for example. The complexity of land-surface and atmospheric processes limits confidence in regional projections of precipitation change, especially over land, although there is a component of a “wet-get-wetter” and “dry-get-drier” response over oceans at the large scale. Even so, there is high confidence that, as the climate warms, extreme precipitation rates (e.g., on daily timescales) will increase faster than the time average. Changes in local extremes on daily and sub-daily timescales are expected to increase by roughly 5 to10% per °C of warming (medium confidence). {7.6, 9.5.4}
For the near- and long-term, CMIP5 projections confirm a clear tendency for increases in heavy precipitation events in the global mean seen in the AR4, but there are substantial variations across regions (TFE.9, Figure 1). Over most of the mid-latitude land-masses and over wet tropical regions, extreme precipitation will very likely be more intense and more frequent in a warmer world. {11.3.2, 12.4.5}
Floods and Droughts
There continues to be a lack of evidence and thus low confidence regarding the sign of trend in the magnitude and/or frequency of floods on a global scale over the instrumental record. There is high confidence that past floods larger than those recorded since 1900 have occurred during the past five centuries in northern and central Europe, western Mediterranean region, and eastern Asia. There is medium confidence that modern large floods are comparable to or surpass historical floods in magnitude and/or frequency in the Near East, India, and central North America. {2.6.2, 5.5.5}
Compelling arguments both for and against significant increases in the land area affected by drought and/or dryness since the mid-20th century have resulted in a low confidence assessment of observed and attributable large-scale trends. This is primarily due to a lack and quality of direct observations, dependencies of inferred trends on the index choice, geographical inconsistencies in the trends and difficulties in distinguishing decadal scale variability from long term trends. On millennial timescales, there is high confidence that proxy information provides evidence of droughts of greater magnitude and longer duration than observed during the 20th century in many regions. There is medium confidence that more megadroughts occurred in monsoon Asia and wetter conditions prevailed in arid Central Asia and the South American monsoon region during the Little Ice Age (1450 to 1850) compared to the Medieval Climate Anomaly (950 to 1250). {2.6.2, 5.5.4, 5.5.5, 10.6.1}
Under the RCP8.5 scenario, projections by the end of the century indicate an increased risk of drought is likely (medium confidence) in presently dry regions linked to regional to global-scale projected decreases in soil moisture. Soil moisture drying is most prominent in the Mediterranean, Southwest US, and southern Africa, consistent with projected changes in the Hadley circulation and increased surface temperatures, and surface drying in these regions is likely (high confidence) by the end of the century under the RCP8.5 scenario. {12.4.5}
Extreme Sea Level
It is likely that the magnitude of extreme high sea level events has increased since 1970 and that most of this rise can be explained by increases in mean sea level. When mean sea level changes is taken into account, changes in extreme high sea levels are reduced to less than 5 mm yr–1 at 94% of tide gauges. In the future it is very likely that there will be a significant increase in the occurrence of sea level extremes and similarly to past observations, this increase will primarily be the result of an increase in mean sea level. {3.7.5, 13.7.2}
Tropical and Extratropical Cyclones
There is low confidence in long-term (centennial) changes in tropical cyclone activity, after accounting for past changes in observing capabilities. However over the satellite era, increases in the frequency and intensity of the strongest storms in the North Atlantic are robust (very high confidence). However, the cause of this increase is debated and there is low confidence in attribution of changes in tropical cyclone activity to human influence due to insufficient observational evidence, lack of physical understanding of the links between anthropogenic drivers of climate and tropical cyclone activity and the low level of agreement between studies as to the relative importance of internal variability, and anthropogenic and natural forcings. {2.6.3, 10.6.1, 14.6.1}
Some high-resolution atmospheric models have realistically simulated tracks and counts of tropical cyclones and models generally are able to capture the general characteristics of storm tracks and extratropical cyclones with evidence of improvement since the AR4. Storm track biases in the North Atlantic have improved slightly, but models still produce a storm track that is too zonal and underestimate cyclone intensity. {9.4.1, 9.4.5}
While projections indicate that it is likely that the global frequency of tropical cyclones will either decrease or remain essentially unchanged, concurrent with a likely increase in both global mean tropical cyclone maximum wind speed and rainfall rates, there is lower confidence in region-specific projections of frequency and intensity. However, due to improvements in model resolution and downscaling techniques, it is more likely than not that the frequency of the most intense storms will increase substantially in the Western North Pacific and North Atlantic under projected 21st century warming (see Figure TS.26). {11.3.2, 14.6.1}
Research subsequent to the AR4 and SREX continues to support a likely poleward shift of storm tracks since the 1950s. However over the last century there is low confidence of a clear trend in storminess due to inconsistencies between studies or lack of long-term data in some parts of the world (particularly in the Southern Hemisphere). {2.6.4, 2.7.6}
Despite systematic biases in simulating storm tracks, most models and studies are in agreement that the global number of extra-tropical cyclones is unlikely to decrease by more than a few per cent. A small poleward shift is likely in the Southern Hemisphere storm track. It is more likely than not (medium confidence) for a projected poleward shift in the North Pacific storm track but it is unlikely that the response of the North Atlantic storm track is a simple poleward shift. There is low confidence in the magnitude of regional storm track changes, and the impact of such changes on regional surface climate. {14.6.2}
Source & ©:
IPCC
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