1
Kelly A. Nugent & H. Damon Matthews (2012). Drivers of Future Northern Latitude Runoff Change,
2
Atmosphere-Ocean, 50:2, 197-206. doi: 10.1080/07055900.2012.658505
3
Abstract
4
Identifying the drivers of changing continental runoff is key to understanding current and
5
predicting future hydrological responses to climate change. Potential drivers of runoff change
6
include changes in precipitation and evaporation due to climate warming, vegetation
7
physiological responses to elevated atmospheric CO2 concentrations, increases in lower-
8
atmospheric aerosols, and anthropogenic land cover change. In this study, we present a series of
9
simulations using an intermediate-complexity climate and carbon cycle model, to assess the
10
contribution of each of these drivers to historical and future continental runoff changes. We
11
present results for global runoff, in addition to northern latitude runoff that discharges into the
12
Arctic and North Atlantic Oceans, so as to identify any potential contribution of increased
13
continental freshwater discharge to changes in North Atlantic deep-water formation. Between
14
1800 and 2100, the model simulated a 26% increase in global runoff, and a 32% runoff increase
15
in the northern latitude region. This increase was driven by a combination of increased
16
precipitation from climate warming, and decreased evapotranspiration due to the physiological
17
response of vegetation to elevated CO2. When isolated, climate warming (and associated
18
changes in precipitation) increased runoff by 16% globally, and by 27% at northern latitudes.
19
Vegetation responses to elevated CO2 led to a 13% increase in global runoff, and a 12% increase
20
in the northern latitude region. These changes in runoff, however, did not affect the strength of
21
the Atlantic Meridional Overturning Circulation, which was affected by surface ocean warming
22
rather than by runoff-induced salinity changes. This study indicates that vegetation physiological
23
responses to elevated CO2 may be contributing to changes in continental runoff at a level similar
24
to that of the direct effect of climate warming.
25
1. Introduction
26
27
Predicting runoff change is challenging given the complexity and uncertainty associated
28
with the climate’s response to anthropogenic forcings and, particularly, the hydrological changes
29
associated with anthropogenic climate warming. However, quantifying a range of possible
30
responses is critical given the important ramifications of a modified hydrological cycle (Allen &
31
Ingram, 2002; Milly, Dunne & Vecchia, 2005). Analysis of historical stream discharge has
32
identified a trend towards increasing continental runoff at high northern latitudes (Labat et al.,
33
2004, Dai et al., 2009). There is some disagreement over historical records at low to mid
34
latitudes, with the result that estimates of global runoff changes range from an increasing trend
35
(Labat et al, 2004) to no statistically detectable trend (Dai et al., 2009). Climate model
36
predictions of future runoff change are also highly variable; though there is general agreement
37
that continental runoff is expected to change on the order of 5-15% per degree of global
38
warming. These projected changes include an anticipated decrease in runoff in many temperate
39
and subtropical regions, and a corresponding increase at high latitudes as a result of an overall
40
intensification of the global hydrological cycle (Solomon et al., 2010).
41
Though these model projections account for changes in precipitation and evaporation
42
associated with anthropogenic climate warming, they generally do not include other potentially
43
important drivers of continental runoff change. In particular, several recent studies have
44
highlighted the potential role of terrestrial vegetation responses to elevated atmospheric CO2
1
45
concentrations as a contributor to global runoff changes. This so-called “physiological forcing”
46
results primarily from decreased stomatal conductance as plants are able to use available soil
47
moisture more efficiently in higher ambient CO2 concentrations. The effect of vegetation
48
physiological forcing has been shown in some regions to amplify the response of runoff to
49
climate warming (Betts et al., 1997; Levis, Foley & Pollard, 2000; Medlyn et al., 2001; Gedney
50
et al., 2006; Betts et al., 2007; Bernacchi et al., 2007; Cao et al., 2009; 2010).
51
In general, vegetation physiological forcing tends to decrease transpiration, leading to
52
increased soil moisture and a corresponding increase in continental runoff (Levis et al., 2000;
53
Piao et al., 2007). This effect can be damped somewhat by plant structural responses to CO2
54
changes whereby enhanced leaf growth can lead to an increase in available surface area for water
55
transpiration and a corresponding increase in moisture recirculation to the atmosphere. In so
56
doing, plant structural responses are thought to at least partially offset the effect of the
57
physiological forcing, though the outcome is also highly dependent on regional landscape
58
characteristics (Betts et al., 1997; Levis et al., 2000; Piao et al., 2007).
59
Anthropogenic land-cover change has also been identified as a potential contributor to
60
historical runoff change through its influence on vegetation and land surface characteristics
61
associated with cropland establishment and abandonment (Piao et al., 2007). For example,
62
deforestation and/or agricultural expansion might be expected to affect runoff as a result of
63
changes in evapotranspiration and soil water retention. In addition, the differing albedos of
64
forests and crops could modify regional radiative forcing and could therefore affect precipitation
65
patterns via altered surface temperatures. At a global scale, however, studies have found
66
relatively little impact of land cover change on historical runoff changes (Gedney et al., 2006).
2
67
Finally, the growing knowledge of the effect of anthropogenic aerosols on climate
68
processes has revealed several potential means by which increased atmospheric aerosols may
69
affect continental runoff. Aerosols in the atmosphere absorb and scatter shortwave radiation
70
resulting in both a cooling of the atmosphere, and a decrease of energy received at Earth’s
71
surface. This has the potential to decrease precipitation (as a result of atmospheric cooling) or to
72
decrease surface evaporation (due to less surface energy availability), either of which could
73
affect overall continental runoff. In addition, anthropogenic aerosols can act as cloud
74
condensation nuclei, leading to changes in their physical and chemical properties, potentially
75
further altering precipitation patterns (Boucher, Myhre & Myhre, 2004; Solomon et al., 2007).
76
Some studies have found a reduction in precipitation and evaporation due to lowered surface
77
solar radiation from aerosol forcing (e.g. Liepert et al., 2004). However, the results of Gedney et
78
al.’s (2006) analysis of 20th century runoff observations showed little evidence of aerosol
79
impacts over northern latitude continents.
80
In this study, we present transient model simulations of historical and future climate
81
change using an intermediate complexity global climate and carbon cycle model. In a series of
82
sensitivity experiments, we assess the relative contributions of physical climate changes due to
83
greenhouse gas increases, vegetation physiological forcing and structural changes, land-use and
84
land-cover change, and anthropogenic aerosols as potential drivers of observed and projected
85
future continental runoff changes. We focus in particular on high northern latitudes, which have
86
seen increased historical runoff, due to the possibility that changes in continental runoff could
87
have implications for ocean circulation and deep water formation in the North Atlantic. Our use
88
of transient simulations with a coupled climate and carbon cycle model enables an examination
3
89
of these potential links between land surface processes and ocean circulation, and in addition,
90
advances previous research focused on equilibrium-only climate simulations.
4
91
2. Methodology
92
93
We use here the University of Victoria’s Earth System Climate Model (UVic ESCM)
94
version 2.9 (Eby et al., 2009, Weaver et al., 2001), a coupled climate-carbon model of
95
intermediate complexity. Version 2.7 of the UVic ESCM was a contributing model to the C4MIP
96
model inter-comparison project (Friedlingstein et al., 2006) and to the Fourth Assessment Report
97
of the Intergovernmental Panel on Climate Change (Meehl et al., 2007). Version 2.9 has an
98
equilibrium climate sensitivity of 3.5 °C and a transient climate response of about 2 °C, which
99
falls approximately within the mid-range of more comprehensive climate models (Meehl et al.,
100
101
2007).
The atmospheric component of the UVic ESCM is a reduced complexity (1 vertical level)
102
energy-moisture balance model, with a spherical grid resolution of 3.6° longitude by 1.8°
103
latitude. The atmosphere model controls heat and moisture transport through Fickian diffusion,
104
as well as moisture transport via advection. Precipitation in the model occurs when relative
105
humidity exceeds 85% (Weaver et al., 2001). This simple atmosphere is coupled to a three-
106
dimensional ocean general circulation model, the Modular Ocean Model version 2.2 (MOM2.2,
107
Pacanowski, 1996) which has 19 vertical levels (Weaver et al, 2001). MOM2.2 is driven by
108
surface boundary conditions from the atmosphere and land surface; in particular, the freshwater
109
flux (as a function of precipitation, evaporation and runoff) is added to the surface ocean in the
110
form of a salinity flux. The maximum Meridional Overturning Circulation (MOC) is also
111
calculated in the ocean model, whose value corresponds to the strength of deep-water formation
112
in the North Atlantic (Weaver et al., 2001).
5
113
The land surface scheme used is the Met Office Surface Exchange Scheme (MOSES)
114
(Cox et al., 1999, Essery et al., 2003), coupled to a biochemical vegetation model (TRIFFID:
115
Top-down Representation of Interactive Foliage and Flora Including Dynamics) which together
116
simulate biophysical processes for five plant functional types found in the model (Cox et al,
117
2001). TRIFFID is a dynamic vegetation model, which simulates the structure and spatial
118
distribution of vegetation in response to changing climate conditions. MOSES simulates land-
119
atmosphere carbon fluxes and calculates various components of the hydrological cycle, including
120
evapotranspiration, soil moisture and runoff. Evapotranspiration in the model is calculated using
121
the Penman-Monteith formulation and varies as a function of both local climate conditions and
122
atmospheric CO2 concentration (Cox et al., 1999, Essery et al., 2003). The version of MOSES
123
implemented in the UVic ESCM is a simplified version of MOSES2 as described in Essery et al
124
(2003); in particular, soil moisture here is stored in a one-meter soil layer, as opposed to the four
125
soil layers in the original version. Excess moisture is diverted to the ocean as runoff via 32
126
realistic river basins, according to a river-routing scheme as described (Weaver et al., 2001). We
127
note that version 2.9 of the UVic ESCM does not represent permafrost, wetlands or frozen soils;
128
this is an area of active model development (Avis et al., 2011), but represents a limitation of the
129
current model in representing Arctic continental runoff responses to climate warming.
130
In our study, we performed five transient model simulations following a 10,000 year
131
spin-up period under pre-industrial forcing. For the reference simulation (ALL) we included all
132
potential drivers of runoff change: the SRES A2 business-as-usual CO2 concentration scenario,
133
including additional prescribed forcing from non-CO2 greenhouse gases (Nakicenovic et al.,
134
2000); physiological and structural responses of vegetation to CO2 changes; historical land-cover
6
135
change from Ramankutty and Foley (1999); and, the direct effect of anthropogenic sulphate
136
aerosols, represented as a prescribed aerosol optical depth following historical and future A2
137
aerosol emissions (see Matthews et al., 2004 and http://climate.uvic.ca/EMICAR5/forcing for
138
more information about aerosol forcing in the UVic model). Subsequently, we performed
139
additional simulations wherein each of the four runoff drivers was turned off. For aerosol effects
140
and land-use change, the effect was removed from each respective simulation by holding the
141
forcing constant at pre-industrial conditions. We isolated the effect of vegetation-CO2
142
physiological forcing by holding CO2 concentrations constant at pre-industrial levels with
143
respect to the vegetation component of the model while allowing CO2 radiative forcing to affect
144
the physical climate model. Conversely, the climate effect of greenhouse gases was isolated by
145
masking the radiative forcing from increased greenhouse gases in the climate model, while
146
allowing CO2 increases alone to affect vegetation growth.
147
Model simulations are summarized in Table 1. In the results presented here, we isolated
148
the effect of individual drivers by calculating the difference between each of the four simulations
149
and the reference simulation ALL. We extracted and analysed runoff changes both globally and
150
for land areas that discharge into the Arctic and North Atlantic Oceans north of 45°N latitude,
151
and between 105°W and 90°E longitude (see grey area marked on Figure 3). Within this region,
152
we calculated runoff, land temperature, precipitation and evaporation over the river basins whose
153
discharge areas were within the specified latitude and longitude ranges. Ocean fields
154
(temperature and salinity) were calculated for all ocean surface areas within the region.
155
156
7
157
3. Results
158
Under present-day conditions, globally-averaged runoff in the model was 1.1 Sv (where 1
159
Sv = 106 kg/s). In our selected northern latitude region, which includes river drainage into the
160
North Atlantic and part of the Arctic Oceans, total runoff was 0.13 Sv at present-day. These
161
values compare well to recent estimates by Dai and Trenberth (2009), who reported a total
162
continental discharge of about 1.15 Sv for the year 2004, and an Arctic Ocean discharge of 0.14
163
Sv.
164
Simulated trends in globally-averaged runoff, global land surface air temperature, land
165
precipitation and land evaporation are shown in Figures 1 (global) and 2 (northern latitude
166
region). Over the 20th century, global runoff in ALL increased by 7.5 mm/year, or 3.4% relative
167
to pre-industrial, accompanying a 0.6 ºC increase in global temperature (Fig. 1A and 1B). This
168
represents a 5.6% increase per degree of global warming, which is consistent with the 4%/ºC
169
reported by Labat et al. (2004) based on historical runoff and temperature observations, though
170
not with Dai and Trenberth (2009) who did not detect a significant trend in global runoff over the
171
past 50 years. Over the specified northern latitude land region, runoff increased by 9.3 mm/yr
172
(Figure 2A), corresponding to a total discharge increase of 0.006 Sv. By way of comparison, Dai
173
and Trenberth (2009) reported a significant increase in continental discharge to the Arctic Ocean
174
of 0.014 Sv between 1948 and 2004, with no significant trend in the Atlantic as a whole; this is
175
broadly consistent with the simulated results in our northern latitude region, which covers a
176
portion of both the Atlantic and Arctic Oceans (see area marked on Figure 3). Over the 21st
177
century, changes in global runoff per degree of global warming were slightly larger (6.3 %/ºC),
178
which is generally consistent with the range of 5-15 %/ºC reported from a range of different
8
179
models by Solomon et al. (2010). In our selected northern latitude region (Fig. 2A), runoff
180
increased by 32% over the 21st century alongside a 4.1 ºC temperature rise – a slightly higher
181
regional sensitivity of 7.8 %/ºC.
182
As can also be seen in Figure 1, greenhouse gas increases (line GG) were the primary
183
driver of global runoff changes as a result of increased precipitation over all land areas at a rate
184
of 1.2 % per degree Celsius of land temperature increase. In the northern latitude region,
185
warming from greenhouse gas increases led to both an increase in precipitation, and a decrease in
186
evaporation, resulting in a relatively larger increase in continental runoff (Figure 2). Vegetation
187
physiological forcing was the second largest contributor to simulated runoff changes. At the
188
global scale, the effect of vegetation physiological forcing was only slightly smaller than the
189
effect of greenhouse gases, though was closer to half the effect of greenhouse gases in the
190
northern latitude region.
191
The mechanism for runoff changes due to physiological forcing was fundamentally
192
different from that of greenhouse gases. In the case of V-CO2, physiological forcing resulted in a
193
large decrease in precipitation, though unlike the case of GG, this was not a function of
194
temperature changes (which were small); rather, precipitation in V-CO2 was driven by a large
195
decrease in evapotranspiration (Fig. 1D and 2D) as a result of increased water-use efficiency at
196
higher CO2. Consequently, there was a clear division between the two mechanisms of
197
influencing changes in continental runoff. Climate warming tended to increase precipitation and
198
drive a parallel increase in runoff (line GG), whereas elevated CO2 tended to induce decreased
199
transpiration, leading to decreased water re-circulation to the atmosphere, increased soil moisture
9
200
and consequent further increase in runoff. The runoff change in the ALL simulation was driven
201
in approximately equal measure by each of these two processes
202
Aerosols (AER) and land cover changes (LCC) were clearly of secondary importance in
203
driving both historical and future runoff changes. Aerosols tended to have the opposite effect as
204
GG, leading to a cooling of land temperatures, and an associated decrease in precipitation and
205
runoff (with little change in evaporation). The effect of aerosols was relatively larger in the
206
northern latitude region, owing to the higher concentration of anthropogenic aerosols over
207
northern continental areas; even here, however, aerosols had less than half the effect of
208
vegetation physiological forcing by the year 2100. Land cover changes had an almost negligible
209
effect on historical runoff changes, and were not relevant to future changes given that areas of
210
anthropogenic land use did not change in the simulations after the year 2000.
211
Land temperatures (Fig. 1B and 2B) increased in both the GG and ALL simulations,
212
though warming in GG was greater (4.8 °C globally and 5.8 °C in the northern region at 2100)
213
than in ALL (4.2 °C globally and 4.8 °C in the northern region) due to the absence of the effect
214
of reflective aerosols in GG. Globally, aerosols accounted for about 0.9 °C cooling at the year
215
2000, with a larger cooling of 1.3 °C in the northern latitude region. Temperatures continued
216
cooling slightly for the first half of the 21st century, then warmed in the latter half of the century
217
due to a decreased aerosol forcing after ~ 2040. V-CO2 and LCC did not play a prominent role in
218
affecting air temperature change over the course of the two centuries. Vegetation structural
219
feedbacks did result in a small amount of warming (~ 0.25 °C globally and in the northern
220
region) as a result of increased leaf area and spatial coverage of vegetation in response to
221
elevated CO2 concentrations, leading to decreased surface albedo. Land-use change led to an
10
222
even smaller amount of cooling (~ 0.1 °C) due to the higher surface albedo associated with
223
agricultural vegetation types compared to forest vegetation.
224
Figure 3 shows the spatial distribution of precipitation minus evaporation (P–E) changes
225
from 1800 to 2100 for each simulation. The spatial patterns seen here reflect the contrasting
226
mechanisms behind simulated runoff changes as discussed in Figure 2. For instance, in GG, a
227
greater P–E increase is seen along coastal regions and in areas of higher elevation; this reflects
228
an increase in precipitation under climate warming that occurs predominantly over the oceans
229
and over orography. Similarly, cooling in AER generated a similar pattern of P–E changes
230
reflecting decreased precipitation. By contrast, P–E changes in V-CO2 are more representative of
231
global vegetation distributions, which reflects the pattern of decreased evapotranspiration driven
232
by vegetation responses to elevated CO2 (CO2 physiological forcing). In general, the spatial
233
pattern of P–E in ALL reflects the combined P–E changes in GG and V-CO2, with smaller
234
contributions from AER and LCC.
235
Finally, we assessed to what extent the simulated runoff increases shown here may have
236
affected deep-water formation in the North Atlantic. Figure 4 shows the simulated change in the
237
Atlantic Meridional Overturning Circulation (AMOC) (Fig. 4A) in response to changes in
238
surface water temperature (Fig. 4B) and salinity (Fig. 4C). For all simulations, AMOC changes
239
appear to correlate with both salinity and temperature variations, with a consistent pattern of
240
decreased (increased) strength of the AMOC associated with increased (decreased) temperature
241
and decreased (increased) salinity. However, it is clear that North Atlantic salinity changes were
242
not driven by changes in continental runoff, but rather by changes in P–E over this ocean region.
243
In particular, the V-CO2 simulation, which showed a large increase in continental runoff, did not
11
244
simulate decreased North Atlantic salinity nor a change in the strength of the AMOC. Instead,
245
ocean circulation appears to be governed primarily by changes in high-latitude temperatures (and
246
the associated change in the latitudinal temperature gradient): climate warming in ALL and GG
247
led to a slowing of the AMOC, whereas climate cooling in AER led to a slightly stronger
248
AMOC. Changes in continental runoff did not have a noticeable effect.
249
250
4. Discussion
251
In this study we investigated the variation in and explanation for future continental runoff
252
changes using transient climate simulations driven by a suite of anthropogenic forcings. Climate
253
change (including climate warming and changes in precipitation) and vegetation physiological
254
responses to elevated CO2 are projected to be the predominant drivers of runoff change for the
255
21st century, contributing in approximately equal measure to simulated global runoff increases.
256
By contrast, land cover change and anthropogenic aerosols had secondary effects on continental
257
runoff trends. We further assessed to what extent simulated changes in continental runoff could
258
affect the strength of North Atlantic deep water formation; while deep water formation did
259
decrease in response to climate change, this occurred as a result of temperature and P–E changes
260
rather than due to runoff-induced salinity changes.
261
The effect of vegetation-CO2 interactions that we have shown here represents the net
262
effect of both physiological and structural responses of vegetation to CO2, though we have also
263
identified the specific climatic outcome of each of these two vegetation processes. The effect of
264
the vegetative structural feedback was primarily to generate a small increase in surface
265
temperature; this occurred as a result of CO2 fertilization, in which forest expansion over
12
266
grasslands led to decreased local surface albedo, consequently increasing absorption of
267
shortwave radiation (see also Matthews, 2007). In contrast, the effect of CO2 physiological
268
forcing was seen predominantly in hydrological cycle changes; elevated CO2 led to increased
269
vegetation water-use efficiency, decreased evapotranspiration and a consequent increase in
270
continental runoff.
271
Our model results contribute to the growing number of studies suggesting that vegetation
272
physiological forcing from increasing atmospheric CO2 is an important contributor to both runoff
273
changes as well as to the overall climate response to CO2 emissions (Betts et al., 2007; Boucher
274
et al., 2009 Cao et al., 2010). The effect of vegetation physiological forcing shown here supports
275
the findings Betts et al. (2007) and Boucher et al. (2009), who use a similar version of the
276
vegetation model employed here, though coupled in their case to the Hadley Centre global
277
climate model. In our study, the effect of physiological forcing on the hydrological cycle was
278
also substantially larger than the direct effect of vegetation changes on surface temperature; we
279
simulated a relatively small temperature increase due to CO2 fertilization of vegetation (less than
280
0.25 °C increase in the northern latitude region and only 0.1 °C increase averaged over the
281
globe). By contrast, other studies have found a somewhat larger effect; for example, Cao et al.
282
(2010) simulated a ~0.42°C increase in global surface temperature due to vegetation
283
physiological responses to elevated CO2. This difference reflects inter-model uncertainty, as well
284
as differences in methodology between our studies; in particular, we have carried out transient
285
model simulations here, whereas Cao et al. (2010) used equilibrium time-slice simulations. This
286
has implications for the timescale of vegetation responses: while hydrological responses to
287
elevated CO2 occur quickly, the structural response and particularly the lateral expansion of
13
288
vegetation is limited by the fairly slow growth of vegetation in the model. Consequently, we
289
suggest that the transient methodology we have adopted provides a more realistic representation
290
of the relative balance of hydrological and temperature responses to CO2-induced vegetation
291
changes.
292
Historical land cover change did not contribute in any large measure to simulated runoff
293
changes in our study. This is consistent with the conclusions of Gedney et al. (2006), who found
294
that land cover change did not have a significant influence on observed continental runoff
295
changes. There is some evidence that deforestation can lead to local-scale changes in runoff due
296
to decreased evapotranspiration and soil water retention (e.g. Costa et al., 2003, Bradshaw et al.,
297
2007); however, these studies have focused on tropical forests, which have experienced a much
298
larger extent of land cover change compared to Northern latitudes. There is also the potential for
299
interactions between future land-use change and vegetation-CO2 responses. For example,
300
agricultural vegetation is likely to respond differently to CO2 changes as compared to natural
301
vegetation types – a phenomenon which we have not explicitly considered here. In addition, we
302
have considered changes in cropland areas only as a representation of land cover change, and
303
have not included pastures as an additional source of land-conversion, which would likely
304
increase slightly the total effect of land-use change shown here.
305
In the case of aerosols, we did simulate decreased precipitation in response to the cooling
306
induced by increased anthropogenic aerosols, which is consistent with other models, as well as
307
the climate response to aerosols resulting from recent volcanic eruptions (Trenberth and Dai,
308
2009). This change in precipitation led to a small decrease in runoff, both globally and in the
309
northern latitude region considered here. We note, however, that we have considered only the
14
310
direct effect of sulphate aerosols, and not a full representation of the possible direct and indirect
311
effect of a range of anthropogenic aerosols; this could also lead to an underestimate of the effect
312
of aerosols on the global hydrological cycle.
313
We note also, that none of the models which have been used thus far to assess the effect
314
of vegetation physiological forcing on continental runoff dynamics include the impacts of
315
permafrost and changes in frozen soils on high-latitude soil water retention in response to climate
316
warming. This is potentially an important additional driver of runoff changes (see e.g. Rawlins
317
et al., 2003, Slater et al., 2007) which would alter the response of high-latitude runoff to
318
continued climate warming. The response of permafrost to climate warming remains a critical
319
open question regarding both the potential for greenhouse gas feedbacks to climate warming, as
320
well as the possible implications for continental runoff changes.
321
We did not find that simulated runoff changes were sufficient to induce a reduction of the
322
strength of North Atlantic deep water formation. Our model did simulate a decrease in the
323
meridional overturning circulation of 4.5 Sv between 1800 and 2100, from a pre-industrial state
324
of 21.5 Sv (about a 20% decrease). This general trend agrees with previous research; for
325
example, Bryan et al. (2006) showed a 22-26% circulation decline per century in response to a
326
1% per year increase in CO2 forcing. This simulated decrease was driven by temperature
327
increases in the North Atlantic (leading to associated changes in precipitation and sea-surface
328
salinity), as well as by a decreased equator to pole temperature gradient brought about by climate
329
warming. By contrast, simulated runoff changes did not affect the strength of the overturning
330
simulation, as evidenced by the lack of AMOC response to runoff changes from vegetation
331
physiological forcing. Our results suggest that vegetation responses to elevated CO2, while
15
332
important for projections of hydrological cycle changes, are not likely to influence projections of
333
the AMOC response to anthropogenic climate warming. However, we note that a significant
334
potential source of freshwater input to the North Atlantic from melting land ice is not included in
335
this study; this has been identified previously as a more important factor influencing salinity in
336
the North Atlantic than stream discharge (Bryan et al., 2006). Clark et al. (2002) found that only
337
an ~0.1 Sv change to the hydrological cycle was sufficient to induce a thermohaline circulation
338
response during the last glaciation, though we note that recent research has questioned the ability
339
of coarse-resolution ocean models to simulate AMOC responses to continental freshwater
340
discharge (Condron and Winsor, 2011). Nevertheless, our findings represent an ~0.04 Sv
341
increase between 1800 and 2100 in freshwater discharge to the North Atlantic, which if
342
combined with significantly increased land ice melt, may be sufficient to affect a more dramatic
343
deep-water formation response than that simulated here.
344
Considering our findings more broadly, they do suggest the potential for major regional
345
alterations in the hydrological cycle, which could have significant implications for both human
346
and ecological communities. One possible positive implication may be that increased runoff
347
could result in increased water availability for agriculture and other human uses. However, such
348
increases will likely be concentrated in areas that already have ample water supply, whereas
349
more arid regions may be negatively impacted by regional declines in precipitation. Ultimately,
350
the combination of greater soil moisture availability and a warmer climate will require adaptation
351
in how humans make use of the water resources available to us in a changing climate.
352
16
353
5. Conclusion
354
Projecting runoff change for the 21st century is an important step in understanding potential
355
future climate change as anthropogenic activities continue to be the dominant driver of changes
356
in the Earth system. Understanding the complexities of the climate system components will help
357
us prevent harmful climate impacts that have begun as a result of unmitigated temperature
358
increases and changes in the spatial patterns of precipitation. This study concludes that plant
359
responses to atmospheric CO2 concentrations will play an important role, in addition to the direct
360
effect of climate warming, in driving future changes in the global hydrological cycle. Vegetation
361
changes will likely also influence regional patterns of temperature changes, though this effect is
362
secondary to the hydrological effects of vegetation physiological responses to elevated CO2. The
363
relationship between increased freshwater discharge and ocean circulation response was shown
364
to be less significant, with changes in North Atlantic deep water formation responding primarily
365
to temperature changes rather than changes in continental runoff. The use of transient
366
simulations with our coupled climate-carbon cycle model has been a useful approach as it has
367
enabled us to investigate potential links between land surface processes and ocean circulation,
368
while advancing previous runoff change research that, thus far, has focused largely on
369
equilibrium-only climate simulations.
17
370
6. References
371
Allen, M. R., & Ingram, W. J. (2002). Constraints on future changes in climate and the
372
373
374
375
hydrologic cycle. Nature, 419(6903), 224-+. doi:10.1038/nature01092
Avis, C. A., Weaver, A. J. & Meissner, K. J. (2011) Reduction in areal extent of high-latitude
wetlands in response to permafrost thaw. Nature Geoscience, Advance Online Publication.
Bernacchi, C. J., Kimball, B. A., Quarles, D. R., Long, S. P., & Ort, D. R. (2007). Decreases in
376
stomatal conductance of soybean under open-air elevation of [CO2] are closely coupled
377
with decreases in ecosystem evapotranspiration. Plant Physiology, 143(1), 134-144.
378
doi:10.1104/pp.106.089557
379
380
Betts, R. A., Cox, P. M., Lee, S. E., & Woodward, F. I. (1997). Contrasting physiological and
structural vegetation feedbacks in climate change simulations. Nature, 387(6635), 796-799.
381
Betts, R. A., Boucher, O., Collins, M., Cox, P. M., Falloon, P. D., Gedney, N., Hemming, D. L.,
382
Huntingford, C., Jones, C. D., Sexton, D. M. H., & Webb, M. J. (2007). Projected increase
383
in continental runoff due to plant responses to increasing carbon dioxide. Nature,
384
448(7157), 1037-U5. doi:10.1038/nature06045
385
Boucher, O., Myhre, G., & Myhre, A. (2004). Direct human influence of irrigation on
386
atmospheric water vapour and climate. Climate Dynamics, 22(6-7), 597-603.
387
doi:10.1007/s00382-004-0402-4
18
388
Boucher, O., Jones, A. & Betts, R.A. (2009). Climate response to the physiological impact of
389
carbon dioxide on plants in the Met Office Unified Model HadCM3. Climate Dynamics,
390
32(2-3), 237-49. doi: 10.1007/s00382-008-0459-6
391
Bradshaw, C. J. A., Sodhi, N. S., Peh, K. S.-H., & Brook, B. R. (2007). Global evidence that
392
deforestation amplifies flood risk and severity in the developing world. Global Change
393
Biology, 13 (11), 2379-2395.
394
Bryan, F. O., Danabasoglu, G., Nakashiki, N., Yoshida, Y., Kim, D. H., Tsutsui, J., & Doney, S.
395
C. (2006). Response of the North Atlantic thermohaline circulation and ventilation to
396
increasing carbon dioxide in CCSM3. Journal of Climate, 19(11), 2382-2397.
397
Cao, L., Bala, G., Caldeira, K., Nemani, R., & Ban-Weiss, G. (2009). Climate response to
398
physiological forcing of carbon dioxide simulated by the coupled community atmosphere
399
model (CAM3.1) and community land model (CLM3.0). Geophysical Research Letters,
400
36, L10402. doi:10.1029/2009GL037724
401
Cao, L., Bala, G., Caldeira, K., Nemani, R., & Ban-Weiss, G. (2010). Importance of carbon
402
dioxide physiological forcing to future climate change. Proceedings of the National
403
Academy of Sciences of the United States of America, 107(21), 9513-9518.
404
doi:10.1073/pnas.0913000107
405
406
Clark, P. U., Pisias, N. G., Stocker, T. F., & Weaver, A. J. (2002). The role of the thermohaline
circulation in abrupt climate change. Nature, 415(6874), 863-869. doi:10.1038/415863a
19
407
408
Condron, A. & Winsor, P. (2011) A subtropical fate awaited freshwater discharged from glacial
Lake Agassiz. Geophysical Research Letters, 38, L03705.
409
Costa, M. H., Botta, A., & Cardille, J. A. (2003) Effects of large-scale changes in land cover on
410
the discharge of the Tocantins River, Southeastern Amazonia. Journal of Hydrology, 283,
411
206-217.
412
Cox, P. M., Betts, R. A., Bunton, C. B., Essery, R. L. H., Rowntree, P. R., & Smith, J. (1999).
413
The impact of new land surface physics on the GCM simulation of climate and climate
414
sensitivity. Climate Dynamics, 15(3), 183-203. doi:10.1007/s003820050276
415
416
417
Dai, A., Qian, T., Trenberth, K. E., & Milliman, J. D. (2009). Changes in continental freshwater
discharge from 1948 to 2004. Journal of Climate, 22(10), 2773-2792.
Eby, M., Zickfield, K., Montenegro, A., Archer, D., Meissner, K.J. & Weaver, A.J. (2009).
418
Lifetime of anthropogenic climate change: millennial time scales of potential CO2 and
419
surface temperature perturbations. Journal of Climate, 22(10), 2501-11.
420
doi:10.1175/2008JCL2554.1
421
Essery, R. L. H., Best, M. J., Betts, R. A., Cox, P. M. & Taylor, C. M. (2003) Explicit
422
Representation of Subgrid Heterogeneity in a GCM Land Surface Scheme. Journal of
423
Hydrometeorology, 4, 530-543.
424
425
Friedlingstein, P., Cox, P., Betts, R., Bopp, L., Von Bloh, W., Brovkin, V., Cadule, P., Doney,
S., Eby, M., Fung, I., Bala, G., John, J., Jones, C., Joos, F., Kato, T., Kawamiya, M.,
20
426
Knorr, W., Lindsay, K., Matthews, H. D., Raddatz, T., Rayner, P., Reick, C., Roeckner, E.,
427
Schnitzler, K. -., Schnur, R., Strassmann, K., Weaver, A. J., Yoshikawa, C., & Zeng, N.
428
(2006). Climate-carbon cycle feedback analysis: Results from the (CMIP)-M-4 model
429
intercomparison. Journal of Climate, 19(14), 3337-3353.
430
Gedney, N., Cox, P. M., Betts, R. A., Boucher, O., Huntingford, C., & Stott, P. A. (2006).
431
Detection of a direct carbon dioxide effect in continental river runoff records. Nature,
432
439(7078), 835-838. doi:10.1038/nature04504
433
Labat, D., Godderis, Y., Probst, J. L., & Guyot, J. L. (2004). Evidence for global runoff increase
434
related to climate warming. Advances in Water Resources, 27(6), 631-642.
435
doi:10.1016/j.advwatres.2004.02.020
436
437
438
Levis, S., Foley, J. A., & Pollard, D. (2000). Large-scale vegetation feedbacks on a doubled CO2
climate. Journal of Climate, 13(7), 1313-1325.
Liepert, B. G., Feichter, J., Lohmann, U., & Roeckner, E. (2004). Can aerosols spin down the
439
water cycle in a warmer and moister world? Geophysical Research Letters, 31(6), L06207.
440
doi:10.1029/2003GL019060
441
442
Matthews, H. D. (2007) Implication of CO2 fertilization for future climate change in a coupled
climate-carbon model. Global Change Biology, 13, 1068-1078.
21
443
Matthews, H. D., Weaver, A. J., Meissner, K. J., Gillett, N. P. and Eby, M. (2004) Natural and
444
anthropogenic climate change: incorporating historical land cover change, vegetation
445
dynamics and the global carbon cycle. Climate Dynamics, 22: 461–479.
446
Medlyn, B. E., Barton, C. V. M., Broadmeadow, M. S. J., Ceulemans, R., De Angelis, P.,
447
Forstreuter, M., Freeman, M., Jackson, S. B., Kellomaki, S., Laitat, E., Rey, A., Roberntz,
448
P., Sigurdsson, B. D., Strassemeyer, J., Wang, K., Curtis, P. S., & Jarvis, P. G. (2001).
449
Stomatal conductance of forest species after long-term exposure to elevated CO2
450
concentration: A synthesis. New Phytologist, 149(2), 247-264.
451
Meehl, G.A., T.F. Stocker, W.D. Collins, P. Friedlingstein, A.T. Gaye, J.M. Gregory, A. Kitoh,
452
R. Knutti, J.M. Murphy, A. Noda, S.C.B. Raper, I.G. Watterson, A.J. Weaver and Z.-C.
453
Zhao, 2007: Global Climate Projections. In: Climate Change 2007: The Physical Science
454
Basis. Contribution of Working Group I to the Fourth Assessment Report of the
455
Intergovernmental Panel on Climate Change [Solomon, S., D. Qin, M. Manning, Z. Chen,
456
M. Marquis, K.B. Averyt, M. Tignor and H.L. Miller (eds.)]. Cambridge University Press,
457
Cambridge, United Kingdom and New York, NY, USA.
458
Milly, P. C. D., Dunne, K. A., & Vecchia, A. V. (2005). Global pattern of trends in streamflow
459
and water availability in a changing climate. Nature, 438(7066), 347-350.
460
doi:10.1038/nature04312
461
Nakicenovic, N., Alcamo, J., Davis, G., de Vries, B., Fenhann, J., Gaffin, S… Dadi, Z. (2000).
462
Special Report on Emissions Scenarios: a special report of working group III of the
22
463
Intergovernmental Panel on Climate Change. Cambridge University Press, Cambridge,
464
U.K., 599 pp.
465
466
467
Pacanowski, R.C. Ed. (1996). MOM 2 version 2.0 (beta): documention, user’s guide and
reference manual. GFDL Ocean Technical Report 3.2, USA, 350 pages.
Piao, S., Friedlingstein, P., Ciais, P., de Noblet-Ducoudre, N., Labat, D., & Zaehle, S. (2007).
468
Changes in climate and land use have a larger direct impact than rising CO2 on global river
469
runoff trends. Proceedings of the National Academy of Sciences of the United States of
470
America, 104(39), 15242-15247. doi:10.1073/pnas.0707213104
471
Ramankutty N., Foley J.A. (1999) Estimating historical changes in land cover: croplands from
472
1700 to 1992. Global Biogeochemical Cycles, 13: 997-1027
473
10.1029/1999GB900046
474
doi:
Rawlins, M. A., Lammers, R. B., Frolkin, S., Fekete, B. M. & Vorosmarty, C. J. (2003)
475
Simulating pan-Arctic runoff with a macro-scale terrestrial water balance model.
476
Hydrological Processes, 17, 2521–2539.
477
Slater, A. G., Bohn, T. J., McCreight, J. L., Serreze, M. C. & Lettenmaier, D. P. (2007) A
478
multimodel simulation of pan-Arctic hydrology. Journal of Geophysical Research, 112,
479
G04S45.
480
481
Solomon, S., D. Qin, M. Manning, R.B. Alley, T. Berntsen, N.L. Bindoff, … D. Wratt (2007)
Technical Summary. In: Climate Change 2007: The Physical Science Basis. Contribution
23
482
of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on
483
Climate Change [Solomon, S., D. Qin, M. Manning, Z. Chen, M. Marquis, K.B. Averyt,
484
M. Tignor and H.L. Miller (Eds.)]. Cambridge, Cambridge University Press.
485
Solomon, S., Battisti, D., Doney, S., Hayhoe, K., Held. I. M., Lettenmaier, D. P., Lobell, D.,
486
Matthews, D., Pierrhumbert, R., Raphael, M., Richels, R., Root, T. L., Steffen, K., Tebaldi,
487
C., & Yohe, G. W. (2010) Climate Stabilization Targets: emissions, concentrations and
488
impacts over decades to millennia. The National Academies Press, Washington, DC, USA.
489
Trenberth, K. E. & Dai, A. (2009) Effects of Mount Pinatubo volcanic eruption on the
490
hydrological cycle as an analog of geoengineering, Geophysical Research Letters, 34,
491
L15702.
492
Trenberth, K.E., P.D. Jones, P. Ambenje, R. Bojariu, D. Easterling, A. Klein Tank, D. Parker, F.
493
Rahimzadeh, J.A. Renwick, M. Rusticucci, B. Soden and P. Zhai, 2007: Observations:
494
Surface and Atmospheric Climate Change. In: Climate Change 2007: The Physical
495
Science Basis. Contribution of Working Group I to the Fourth Assessment Report of the
496
Intergovernmental Panel on Climate Change [Solomon, S., D. Qin, M. Manning, Z. Chen,
497
M. Marquis, K.B. Averyt, M. Tignor and H.L. Miller (eds.)]. Cambridge University Press,
498
Cambridge, United Kingdom and New York, NY, USA.
499
Weaver, A. J., Eby, M., Wiebe, E. C., Bitz, C. M., Duffy, P. B., Ewen, T. L., Fanning, A. F.,
500
Holland, M. M., MacFadyen, A., Matthews, H. D., Meissner, K. J., Saenko, O.,
501
Schmittner, A., Wang, H. X., & Yoshimori, M. (2001). The UVic earth system climate
24
502
model: Model description, climatology, and applications to past, present and future
503
climates. Atmosphere-Ocean, 39(4), 361-428.
504
25
505
7. Figure Legends
506
Figure 1:
507
Simulated global trend in A) continental runoff, B) land temperature, C) precipitation over land,
508
and D) evaporation over land, in response to the drivers ALL, GG, V- CO2 , LCC and AER.
509
Figure 2:
510
Simulated northern latitude trend in A) continental runoff, B) land temperature, C) precipitation
511
over land, and D) evaporation over land, in response to the drivers ALL, GG, V- CO2 , LCC and
512
AER. The region shown here covers all river basins that drain into the North Atlantic and Arctic
513
Oceans north of 45 °N latitude, and between 105°W and 90°E longitude, as indicated by the
514
shaded box in Figure 3.
515
Figure 3:
516
Spatial distribution of precipitation – evapotranspiration (P–E) change between 1800 and 2100,
517
in response to each driver: A) ALL simulation; B) greenhouse gas forcing; C) vegetation
518
physiological forcing; E) aerosol forcing; E) land cover change. The grey bounding box
519
represents our selected northern latitude region.
520
Figure 4:
521
Simulated trends in the Atlantic Meridional Overturning Circulation (AMOC) (A), surface ocean
522
temperature changes (B) and surface ocean salinity changes (C) in response to the drivers ALL,
523
GG, V- CO2 , LCC and AER. Surface ocean temperature and salinity values are averaged over
524
the North Atlantic and Arctic Oceans, north of 45 °N latitude, and between 105°W and 90°E
525
longitude.
26
526
Table 1: Description of model simulations.
Simulation
ALL
Reference simulation with all runoff drivers included: CO2+non-CO2 radiative
forcing; Vegetation-CO2 physiological forcing; direct effect of sulphate aerosols
and historical land cover change.
GG
Effect of CO2 and non-CO2 greenhouse gas forcing alone*
V-CO2
Effect of vegetation physiological forcing alone*
AER
Effect of direct sulphate aerosol forcing alone*
LCC
Effect of historical land cover change alone*
* Effect of single driver calculated as the difference between a simulation without this process, and the reference simulation ALL.
527
528
27
529
530
28
531
532
29
533
30
534
31