Thursday, December 22, 2011

1180342271.txt

From: Gavin Schmidt <gschmidt@giss.nasa.gov>
To: Phil Jones <p.jones@uea.ac.uk>
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 <ammann@ucar.edu>

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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.

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