From: Gavin Schmidt To: Phil Jones Subject: Wengen section Date: Mon, 28 May 2007 04:51:11 -0400 (EDT) Reply-to: gschmidt@giss.nasa.gov Cc: mann@psu.edu, Caspar Ammann Hi Phil, sorry for the long delay. But here is a first draft of the forcings and models section I was supposed to take the lead on. Hopefully, we can merge that with whatever Caspar has. Thanks Gavin ================ 4 Forcing (GS/CA/EZ) 4-5pp Histories (CA) How models see the forcings, especially wrt aerosols/ozone and increasing model complexities (GS) An important reason for improving climate reconstructions of the past few millenia is that these reconstructions can help us both evaluate climate model responses and sharpen our understanding of important mechanisms and feedbacks. Therefore, a parallel task to improving climate reconstructions is to assess and independently constrain forcings on the climate system over that period. Forcings can generically be described as external effects on a specific system. Responses within that system that also themselves have an impact on its internal state are described as feeebacks. For the atmosphere, sea surface temperature changes could therefore be considered a forcing, but in a coupled ocean-atmosphere model they could be a feedback to another external factor or be intrinsic to the coupled system. Thus the distinction between forcings and feedbacks is not defined a priori, but is a function of the scope of the modelled system. This becomes especially important when dealing with the bio-geo-chemical processes in climate that effect the trace gas concentrations (CO2 and CH4) or aerosols. For example, if a model contains a carbon cycle, than the CO2 variations as a function of climate will be a feedback, but for a simpler physical model, CO2 is often imposed directly as a forcing from observations, regardless of whether in the real world it was a feedback to another change, or a result of human industrial activity. It is useful to consider the pre-industrial period (pre-1850 or so) seperately from the more recent past, since the human influence on many aspects of atmospheric composition has increased dramatically in the 20th Century. In particular, aerosol and land use changes are poorly constrained prior to the late 20th Century and have large uncertainties. Note however, there may conceivably be a role for human activities even prior to the 19th Century due to early argiculatural activity (Ruddiman, 2003; Goosse et al, 2005). In pre-industrial periods, forcings can be usefully separated into purely external changes (variations of solar activity, volcanic eruptions, orbital variation), and those which are intrinsic to the Earth system (greenhouse gases, aerosols, vegetation etc.). Those changes in Earth system elements will occur predominantly as feedbacks to other changes (whether externally forced or simply as a function of internal climate 'noise'). In the more recent past, the human role in affecting atmospheric composition (trace gases and aerosols) and land use have dominated over natural processes and so these changes can, to large extent, be considered external forcings as well. Traditionally, the 'system' that is most usually implied when talking about forcings and feedbacks are the 'fast' components atmosphere-land surface-upper ocean system that, not coincidentally, corresponds to the physics contained within atmospheric general circulation models (AGCMs) coupled to a slab ocean. What is not included (and therefore considered as a forcing according to our previous definition) are 'slow' changes in vegetation, ice sheets or the carbon cycle. In the real world these features will change as a function of other climate changes, and in fact may do so on relatively 'fast' (i..e multi-decadal) timescales. Our choice then of the appropriate 'climate system' is thus slightly arbitrary and does not give a complete picture of the long term sensitivity of the real climate. These distinctions become important because the records available for atmospheric composition do not record the distinction between feedback or forcing, they simply give, for instance, the history of CO2 and CH4. Depending on the modelled system, those records will either be a modelling input, or a modelling target. While there are good records for some factors (particularly the well mixed greenhouse gases such as CO2 and CH4), records for others are either hopelessly incomplete (dust, vegetation) due to poor spatial or temporal resolution or non-existant (e.g. ozone). Thus estimates of the magnitude of these forcings can only be made using a model-based approach. This can be done using GCMs that include more Earth system components (interactive aerosols, chemistry, dynamic vegetation, carbon cycles etc.), but these models are still very much a work in progress and have not been used extensively for paleo-climatic purposes. Some initial attempts have been made for select feedbacks and forcings (Gerber et al, 2003; Goosse et al 2006) but a comprehensive assessment over the millennia prior to the pre-industrial does not yet exist. Even for those forcings for which good records exist, there is a question of they are represented within the models. This is not so much of an issue for the well-mixed greenhouse gases (CO2, N2O, CH4) since there is a sophisticated literature and history of including them within models (IPCC, 2001) though some aspects, such as minor short-wave absorption effects for CH4 and N2O are still not universally included (Collins et al, 2006). However, solar effects have been treated in quite varied ways. The most straightforward way of including solar irradiance effects on climate is to change the solar 'constant' (preferably described as total solar irradiance - TSI). However, observations show that solar variability is highly dependent on wavelength with UV bands having about 10 times as much amplitude of change than TSI over a solar cycle (Lean, 2000). Thus including this spectral variation for all solar changes allows for a slightly different behaviour (larger solar-induced changes in the stratosphere where the UV is mostly absorbed for instance). Additionally, the changes in UV affect ozone production in both the stratosphere and troposphere, and this mechanism has been shown to affect both the total radiative forcing and dynamical responses (Haigh 1996, Shindell et al 2001; 2006). Within a chemistry climate model this effect would potentially modify the radiative impact of the original solar forcing, but could also be included as an additional (parameterised) forcing in standard GCMs. There is also a potential effect from the indirect effect of solar magnetic variability on the sheilding of cosmic rays, which have been theorised to affect the production of cloud condensation nuclei (Dickinson, 1975). However, there have been no quantitative calculations of the magnitude of this effect (which would require a full study of the relevant aerosol and cloud microphysics), and so its impact on climate is not (yet) been included. Large volcanic eruptions produce significant amounts of sulpher dioxide (SO2). If this is injected into the tropical stratosphere during a particularly explosive eruption, the resulting sulphate can persist in the atmosphere for a number of years (e.g. Pinatubo in 1991). Less explosive, but more persistent eruptions (e.g. Laki in 1789??) can still affect climate though in a more regional way and for a shorter term (Oman et al, 2005). These aerosols have both a shortwave (reflective) and longwave (absorbing) impact on the radiation and their local impact on stratospheric heating can have important dynamical effects. It is therefore better to include the aerosol absorber directly in the radiative transfer code. However, in less sophisticated models, the impact of the aerosols has been parameterised as the equivalent decrease in TSI. For extreme eruptions it has been hypothesised that sulphate production might saturate the oxidative capacity of the stratosphere leaving significant amounts of residual SO2. This gas is a greenhouse gas and would have an opposite effect to the cooling aerosols. This effect however has not yet been quantified. Land cover changes have occured both due to deliberate modification by humans (deforestation, imposed fire regimes, arguculture) as well as a feedback to climate change (the desertification of the Sahara ca. 5500 yrs ago). Changing vegetation in a standard model affects the seasonal cycle of albedo, the surface roughness, the impact of snow, evapotranspiration (through different rooting depths) etc. However, modelling of the yearly cycle of crops, or incorporating the effects of large scale irrigation are still very much a work in progress. Aerosol changes over the last few milllenia are very poorly constrained (if at all). These might have arisen from climatically or human driven changes in dust emissions, ocean biology feedbacks on circulation change, or climate impacts on the emission volatile organics from plants (which also have an impact on ozone chemistry). Some work on modelling a subset of those effects has been done for the last glacial maximum or the 8.2 kyr event (LeGrande et al, 2006), but there have been no quantitative estimates for the late Holocene (prior to the industrial period). Due to the relative expense of doing millennial simulations with state-of-the-art GCMs, exisiting simulations have generally done the minimum required to include relevant solar, GHG and volcanic forcings. Progress can be expected relatively soon on more sophisticated treatments of those forcings and the first quantitative estimates of additional effects. ============= *--------------------------------------------------------------------* | Gavin Schmidt NASA/Goddard Institute for Space Studies | | 2880 Broadway | | Tel: (212) 678 5627 New York, NY 10025 | | | | gschmidt@giss.nasa.gov http://www.giss.nasa.gov/~gavin | *--------------------------------------------------------------------*