Download Radiative forcing of climate by historical land cover change

GEOPHYSICAL RESEARCH LETTERS, VOL. 30, NO. 2, 1055, doi:10.1029/2002GL016098, 2003
Radiative forcing of climate by historical land cover change
H. Damon Matthews, Andrew J. Weaver, Michael Eby, and Katrin J. Meissner
School of Earth and Ocean Sciences, University of Victoria, Victoria, British Columbia, Canada
Received 13 August 2002; revised 5 September 2002; accepted 25 September 2002; published 22 January 2003.
[1] The radiative effect of changing human land-use
patterns on the climate of the past 300 years is discussed
through analysis of a series of equilibrium and transient
climate simulations using the UVic Earth System Climate
Model. Land-surface changes are prescribed through
varying land cover type, representing the replacement of
natural vegetation by human agricultural systems from
1700 to 1992. All land cover simulations show a cooling
in the range of 0.09 to 0.22!C with larger regional changes
INDEX TERMS: 3322
caused by local positive feedbacks.
Meteorology and Atmospheric Dynamics: Land/atmosphere
interactions. Citation: Matthews, H. D., A. J. Weaver, M.
Eby, and K. J. Meissner, Radiative forcing of climate by
historical land cover change, Geophys. Res. Lett., 30(2), 1055,
doi:10.1029/2002GL016098, 2003.
1. Introduction
[2] In the study of anthropogenic climate change, there is
a growing interest in the ways that human activities interact
to affect the climate system. In particular, human land-use
over the last several centuries, through large scale changes
to the global vegetation cover, has had a potentially significant effect on global climate. These land cover changes
have recently been quantified in global datasets that provide
estimates of historical land cover change over the past three
centuries [Klein Goldewijk, 2001; Ramankutty and Foley,
1999]. These datasets now allow for historically realistic
analyses of land cover change using general circulation
climate models.
[3] Land cover change can influence global climate in a
number of ways, by altering parameters such as leaf area
index, root distribution, soil water holding capacity, surface
resistance to evaporation, fractional vegetation cover, surface albedo, and fluxes and storage of carbon and other
nutrients [see for example Betts, 2001; Bolin et al., 2000].
As many of these changes have quite complex effects, this
paper focuses on the effects of two key processes, whose
mechanisms are conceptually simple. First, as natural vegetation is replaced by cropland, surface albedo generally
increases, as trees are replaced by agricultural crops. Second, the energy available for evaporation is affected and the
plant transpiration pathway is altered as a result of a
modified vegetation cover.
[4] In a review of contemporary climate forcings, Hansen
et al. [1998] estimate the radiative forcing resulting from
land-use induced changes to surface albedo to be !0.2 ± 0.2
W/m2. In addition, several recent observational and modelling studies have noted cooling on both local and global
scales resulting from scenarios of land cover change [see
Copyright 2003 by the American Geophysical Union.
0094-8276/03/2002GL016098$05.00
Betts, 2001; Bonan, 2001; Brovkin et al., 1999; Chase et al.,
2001; Govindasamy et al., 2001; Zhao et al., 2001, and
others cited therein].
[5] This paper seeks to quantify the global temperature
change resulting from the radiative forcing associated with
historical land cover change by incorporating a global
croplands dataset into an intermediate complexity climate
model. Results from a series of equilibrium and transient
climate runs are presented so as to compare the climate
response to land cover change over the past 300 years with
the responses to greenhouse gas and solar orbital forcing.
This work uses the UVic Earth System Model [Weaver et
al., 2001], which in contrast to previous modelling studies,
includes a full ocean general circulation model (GCM) and
a thermodynamic/dynamic sea-ice model. As such, this
study allows for an estimate of the radiative impact of
historical land cover change on global climate, including
feedbacks associated with ocean circulation and changing
sea-ice distribution.
2. Methodology
2.1. Model
[6] The UVic Earth System Model [Weaver et al., 2001]
is an intermediate complexity climate model comprising a
full ocean GCM (version 2.2 of the GFDL Modular Ocean
Model) coupled to a vertically integrated energy/moisture
balance atmospheric model and a dynamic/thermodynamic
sea-ice model. Surface wind stress as well as verticallyintegrated atmospheric winds used for advection of moisture are specified from NCEP reanalysis data. A dynamic
wind feedback parameterisation allows for wind perturbations to be applied when simulating past climates. The
coupled ocean-atmosphere model has a resolution of 3.6!
in longitude and 1.8! in latitude, and conserves both
energy and water without the use of flux adjustments
[Weaver et al., 2001].
[7] In this work, the UVic Model is further coupled to a
simple land surface model. The model used is a version of
Manabe’s [1969] simple bucket model, modified to allow
for on-line calculation of the aerodynamic resistance to
evaporation; the Dalton number for evaporation over land
is calculated from a spatially variable surface roughness
length, following the methodology of [Brutsaert, 1982]. A
single dominant vegetation type is specified at each grid
cell, using seven vegetation types based on the vegetation
field of DeFries and Townsend [1994]. The specified
vegetation type determines the surface albedo, as well as
the roughness length used in the calculation of the surface
resistance to evaporation. These values (shown in Table 1)
are obtained by spatially averaging the surface albedo and
roughness length fields provided in Sellers et al. [1996]
according to the chosen vegetation type categories. The
27 - 1
27 - 2
MATTHEWS ET AL.: LAND COVER CHANGE AND CLIMATE
Table 1. Vegetation Types and Specified Parameter Values
(Derived From DeFries and Townsend [1994] and Sellers et al.
[1996])
Vegetation Type
Albedo
Roughness length (m)
Tropical Forest
Temperate/Boreal Forest
Grassland/Savanna/Cropland
Shrubland
Tundra
Desert
Rock/Ice
0.13
0.11
0.17
0.17
0.20
0.28
0.14
2.86
0.91
0.11
0.05
0.04
0.04
0.02
presence of snow or sea-ice increases the surface albedo
locally by 0.18.
2.2. Experiments
[8] Our model results are from three 2000 year equilibrium runs and four transient model runs covering the
period from 1700 to 1992. The croplands dataset used is
that of Ramankutty and Foley [1999], which provides a
potential or natural vegetation field (representing vegetation
distributions in the absence of human disturbance) and
fractional cropland areas at each grid cell yearly from
1700 to 1992. As Ramankutty and Foley [1999] consider
only historical changes in cropland area, and not other land
cover changes such as pasture and extensively logged areas
[see also Klein Goldewijk, 2001], this dataset could be
expected to generate a lower bound on the estimate of the
radiative effect of land cover changes.
[9] The three equilibrium runs presented here are differentiated by the specified vegetation fields. Equilibrium (1)
uses present day vegetation as determined from satellite data
[DeFries and Townsend, 1994]. Equilibria (2) and (3) use
the fractional cropland areas of Ramankutty and Foley
[1999] corresponding to years 1992 and 1700, respectively
(as shown in Figure 1), superimposed onto the potential
vegetation field. All three equilibria use present day CO2
concentration (365 ppm) and orbital parameters. As the
croplands dataset used here does not include all land cover
changes that have occurred since 1700, it would be expected
that equilibrium 1 (present day vegetation) will show a
stronger land cover signal than equilibrium 2 (1992 land
cover). Differences in the definition and placement of
vegetation types between the two vegetation fields does
have the potential to have a confounding influence on the
land cover signal in equilibrium (1), but this is nevertheless
a useful experiment to estimate the effect of land cover
changes not captured in the Ramankutty and Foley [1999]
dataset.
[10] Transient climate simulations begin from a control
equilibrium climate defined by year 1700 land cover (as in
equilibrium 3 above), pre-industrial CO2 (280 ppm) and
year 1700 orbital parameters. This control equilibrium is
spun up for 1700 years, and transient simulations are run
from year 1700 to present. The three transient runs are
forced by: (1) changing croplands (land cover) only; (2)
changing CO2 (as represented by the observed increase to
present-day levels beginning in 1850); and (3) changing
land cover and atmospheric CO2. A control experiment (4)
is also conducted with changing orbital parameters, but not
other radiative forcings. In runs 1 and 3, croplands are
specified yearly from 1700 to 1992, and surface albedos and
roughness lengths are changed according to the fractional
cropland area in each grid cell.
3. Results
3.1. Equilibrium Results
[11] Results of the equilibrium model runs are presented
as differences: (a) present day vegetation equilibrium minus
year 1700 land cover equilibrium; and (b) year 1992 land
cover equilibrium minus year 1700 land cover equilibrium.
Negative differences thus represent a decrease resulting
from changes in land cover between the year 1700 and
the present day.
[12] Figure 2 shows zonally averaged temperature, precipitation and evaporation differences for comparisons (a)
and (b). Both comparisons show a cooling and a decrease in
precipitation at all latitudes. Comparison (a) shows a
globally averaged cooling of 0.22!C and an average precipitation decrease of 14.2 mm/year. While the spatial
pattern is similar, the magnitude of the change shown for
comparison (b) is less: a cooling of 0.10!C and precipitation
decrease of 3.3 mm/year. As can be seen in the shape of the
temperature plot, cooling is amplified in the Northern
Hemisphere, consistent with the location of the majority
of the specified land cover changes (see Figure 1). There is
also an amplification at around 70!S, resulting from local
sea-ice albedo feedbacks.
[13] Much of the decrease in precipitation can be attributed to atmospheric temperature change. However, the large
decreases seen in the tropics and sub-tropics can be further
explained by examining the roughness lengths listed in
Table 1. According to the parameterisation of evaporation
in the bucket model, a larger roughness length results in a
smaller aerodymanic resistance to evaporation and hence a
larger Dalton number and (given equal moisture availability) more evaporation. By changing roughness length alone,
a much larger decrease in evaporation would be expected as
Figure 1. Fractional cropland areas for year 1700 (top)
and 1992 (bottom), from Ramankutty and Foley [1999].
27 - 3
MATTHEWS ET AL.: LAND COVER CHANGE AND CLIMATE
Figure 4. Globally averaged temperature for four transient
runs. Total temperature changes: (1) land cover only:
!0.09!; (2) CO2: +0.76!; (3) land cover +CO2: +0.68!; and
(4) Control: +0.003!.
Figure 2. Zonally averaged equilibrium temperature (top),
precipitation (red lines, bottom) and evaporation (green
lines, bottom) differences for (a) present day vegetation –
1700 land cover and (b) 1992 land cover – 1700 land cover.
a result of a conversion from forest to grassland in the
tropics than by a similar change in the temperate/boreal
regions. As such, the greatest evaporation (and hence
precipitation) differences are seen in the tropics where
roughness length changes are largest, with a less pronounced decrease seen in the northern mid-latitudes.
[14] Spatially-varying temperature differences are shown
in Figure 3. While the majority of the temperature change
can be seen in the Northern Hemisphere, the largest local
changes are the result of local snow and sea-ice albedo
feedbacks. Particularly in the case of comparison (a), these
positive feedbacks serve to amplify the cooling signal
initiated by land cover change. Comparison (b) shows less
cooling, but as in the case of the zonally averaged results,
the spatial pattern of cooling is similar to (a). As the model
used here does not contain atmospheric variability [Weaver
et al., 2001], the differences shown here represent significant changes in the climate mean state.
3.2. Transient Results
[15] Transient runs were chosen so as to compare the
importance of land cover forcing of climate to greenhouse
gas forcing. The results of these runs are shown in Figure 4.
The transient effect of land cover change is !0.09!C, quite
close to that found in comparison (b) in the equilibrium
results (!0.10!C). This indicates that there is little evidence
of an oceanic cooling commitment associated with land
cover change, as the majority of the temperature change
seen in the equilibrium comparison is also seen in the
transient run. While the land cover effect on global temperature is smaller than that of CO2 (+0.76!C), it is of
significant magnitude, and has a noticeable dampening
effect on the total warming when both forcings are combined (+0.68!C). It is also clear that land cover change is
much larger than any model variability or orbitally induced
response seen in the control run (+0.003!C).
[16] Transient changes in global temperature, evaporation, precipitation, and short-wave radiation for the four
runs (as well as a fifth run described below) are summarized
in Table 2. The globally averaged short-wave radiative
forcing resulting from land cover change is !0.15 W/m2,
slightly less than, but within the error bounds of that found
by Hansen et al. [1998].
3.3. Model Sensitivity
[17] As shown in Table 1, the albedo value assigned to
croplands in this model is 0.17. As a globally and tempoTable 2. Total Changes in Temperature, Evaporation, Precipitation and Downward Short-Wave Radiation Absorbed at the Surface
Transient Model Run
Figure 3. Two dimensional temperature differences for
equilibrium comparisons (a) present day vegetation – 1700
land cover (upper panel) and (b) 1992 land cover – 1700
land cover (lower panel).
1.
2.
3.
4.
5.
Land Cover (LC) only
CO2 only
LC + CO2
Control
LC (crop albedo = 0.2)
T (!C) E (mm/yr) P (mm/yr) #SW (W/m2)
!0.09
+0.76
+0.68
+0.003
!0.17
!3.3
+12.1
+9.1
"0
!4.9
!3.3
+11.9
+9.0
"0
!4.9
!0.15
+0.06
!0.09
+0.002
!0.28
27 - 4
MATTHEWS ET AL.: LAND COVER CHANGE AND CLIMATE
rally averaged albedo value for grassland, savanna and
cropland vegetation types, this value reproduces well the
satellite albedo data provided by Sellers et al. [1996].
Previous studies of vegetation albedo, however, have placed
the albedo for agricultural land in the range of 0.12 to 0.25,
depending on the crop type, with only rice having a crop
albedo of less than 0.17 [Wilson and Henderson-Sellers,
1985]. It is not clear that 0.17 is a good representative
albedo for a pure cropland vegetation type. In order to
address this concern, a second land cover transient run was
performed using a cropland albedo value of 0.2 for all
fractional cropland areas applied onto the background
vegetation field. This change almost doubles the albedo
forcing associated with land cover change, as not only do
forest grid cells now see a 50 to 75% larger albedo increase,
but the albedo of grassland, savanna and shrubland grid
cells is also increased from 0.17 to 0.2 in those areas where
croplands are applied.
[18] As can be seen in the last line of Table 2, a higher
specified cropland albedo value has a large impact on the
total temperature change for the transient run. Instead of a
cooling of 0.09!C as in the case of the original run, this
second land cover transient run shows a much larger cooling
of 0.17!C. Based on this result, it is clear that the specified
cropland albedo is of primary importance for determining
the magnitude of the global temperature change forced by
land cover change.
4. Conclusions and Future Directions
[19] This study demonstrates that historical land cover
change has had a significant radiative effect on global
climate over the past three centuries. All model runs show
a cooling associated with land cover change, both in global
averages, and as amplified locally by positive feedbacks. In
the equilibrium runs, global cooling was in the range of
0.10!C to 0.22!C, depending on which vegetation field was
used to represent ‘‘present day’’ vegetation. In the transient
case, the cooling was in the range of 0.09!C to 0.17!C,
depending on the specified albedo value for cropland. From
this analysis, it is clear that uncertainties in surface albedo
values can have a large impact on the magnitude of the
modelled land cover effect. It is also evident from comparing the land cover only transient run (0.09!C cooling) with
equilibrium runs (2) and (3) (0.10!C cooling) that the effect
of land cover change does not carry with it a large oceanic
cooling commitment. This is in contrast to the effect of
greenhouse gas forcing where the large thermal memory of
the ocean has resulted in differences as large as 0.61!C in
the UVic model between transient and equilibrium simulations of the present day climate [Weaver et al., 2001].
[20] This work, and others in this area of research, raises
important questions about the future of human-induced
climate change over the next several centuries. In addition
to the physical changes addressed in this paper, land cover
change has a large impact on the climate system through the
release of stored carbon reserves to the atmosphere [Bolin et
al., 2000]. In the transient climate runs presented here, the
terrestrial carbon cycle was not treated as an interactive
component of the climate system, and it was assumed that
releases of greenhouse gases due to land cover change were
captured by the specified atmospheric CO2. Much research
remains to be done to determine the net effect of land cover
change on climate. In a simulation of Northern Hemisphere
afforestation, for example, [Betts, 2000] found that in some
areas of boreal forest, the warming effect of albedo change
was equivalent to the cooling effect of carbon sequestration,
resulting in a neutral total forcing. This question has yet to
be addressed using realistic scenarios of transient land cover
change.
[21] A related question is that of the effect of land cover
change on the future of the terrestrial carbon sink. In many
instances, land-use change reduces the total ability of the
earth’s terrestrial ecosystems to draw carbon from the
atmosphere [Bolin et al., 2000]. How will continued land
cover change affect the capacity of the terrestrial carbon
sink to respond to and mitigate future climate change?
Questions such as these must remain a focus of research
if we are to continue to improve the accuracy of climate
predictions for the next century.
[22] Acknowledgments. The authors wish to thank N. Ramankutty
and J. Foley for making available the historical land cover dataset used in
this study, as well as M. Claussen, V. Brovkin and two anonymous
reviewers for useful comments and suggestions. Financial support for this
research from NSERC and CFCAS is gratefully acknowledged.
References
Betts, R. A., Offset of the potential corbon sink from boreal forestation by
decreases in surface albedo, Nature, 408, 187 – 190, 2000.
Betts, R. A., Biogeophysical impacts of land use on present-day climate:
near surface temperature and radiative forcing, Atmos. Sci. Lett., 1,
doi:10.1006asle.2001.0023, 2001.
Bolin, B., et al., Global perspective, in Land Use, Land-Use Change, And
Forestry: A Special Report of the Intergovernmental Panel on Climate
Change, edited by R. T. Watson et al., pp. 23 – 51, Camb. Univ. Press, 2000.
Bonan, G. B., Observational evidence for reduction of daily temperature by
croplands in the midwest United States, J. Clim., 14, 2430 – 2442, 2001.
Brovkin, V., et al., Modelling climate response to historical land cover
change, Glob. Ecol. Biogeog., 8, 509 – 517, 1999.
Brutsaert, W., Evaporation into the Atmosphere, D. Reidel, Boston, 1982.
Chase, T., et al., Relative climatic effects of landcover change and elevated
carbon dioxide combined with aerosols: A comparison of model results
and observations, J. Geophys. Res., 106, 31,685 – 31,691, 2001.
DeFries, R., and J. Townsend, NDVI-derived land cover classification at
global scales, Int. J. Rem. Sen., 15, 3567 – 3586, 1994.
Govindasamy, B., P. Duffy, and K. Caldeira, Land use change and Northern
Hemisphere cooling, Geophys. Res. Lett., 28, 291 – 294, 2001.
Hansen, J. E., et al., Climate forcings in the industrial era, Proc. Nat. Acad.
Sci.: USA, 95, 12,753 – 12,758, 1998.
Klein Goldewijk, K., Estimating global land use change over the past 300
years: the HYDE database, Glob. Biogeochem. Cycl., 15, 415 – 433,
2001.
Manabe, S., Climate and the ocean circulation 1. the atmospheric circulation and the hydrology of the earth’s surface, Mon. Wea. Rev., 97, 739 –
774, 1969.
Ramankutty, N., and J. A. Foley, Estimating historical changes in land
cover: croplands from 1700 to 1992, Glob. Biogeochem. Cycl., 13,
997 – 1027, 1999.
Sellers, P., et al., The ISLSCP initiative I global datasets: Surface boundary
conditions and atmospheric forcings for land-atmosphere studies, Bull.
Am. Met. Soc., 77, 1987 – 2005, 1996.
Weaver, A. J., et al., The UVic Earth System Climate Model: Model description, climatology and applications to past, present and future climates, Atmos.-Ocean, 39, 361 – 428, 2001.
Wilson, M., and A. Henderson-Sellers, A global archive of land cover and
soils data for use in general circulation climate models, J. Climatol., 5,
119 – 143, 1985.
Zhao, M., A. Pitman, and C. Chase, The impact of land cover change on the
atmospheric circulation, Clim. Dyn., 17, 467 – 477, 2001.
!!!!!!!!!!!!!!!!!!!!!!
H. D. Matthews, A. J. Weaver, M. Eby, and K. J. Meissner, School of
Earth and Ocean Sciences, University of Victoria, P.O. Box 3055, Victoria,
BC V8W 3P6, Canada. ([email protected])