Centennial Changes of the Global Water Cycle in CMIP5
Models
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Levang, Samuel J., and Raymond W. Schmitt. “Centennial
Changes of the Global Water Cycle in CMIP5 Models.” J.
Climate 28, no. 16 (August 2015): 6489–6502. © 2015 American
Meteorological Society
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http://dx.doi.org/10.1175/jcli-d-15-0143.1
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15 AUGUST 2015
6489
LEVANG AND SCHMITT
Centennial Changes of the Global Water Cycle in CMIP5 Models
SAMUEL J. LEVANG
Massachusetts Institute of Technology–Woods Hole Oceanographic Institution Joint Program in
Oceanography, Woods Hole, Massachusetts
RAYMOND W. SCHMITT
Woods Hole Oceanographic Institution, Woods Hole, Massachusetts
(Manuscript received 19 February 2015, in final form 12 May 2015)
ABSTRACT
The global water cycle is predicted to intensify under various greenhouse gas emissions scenarios. Here the
nature and strength of the expected changes for the ocean in the coming century are assessed by examining the
output of several CMIP5 model runs for the periods 1990–2000 and 2090–2100 and comparing them to a
dataset built from modern observations. Key elements of the water cycle, such as the atmospheric vapor
transport, the evaporation minus precipitation over the ocean, and the surface salinity, show significant
changes over the coming century. The intensification of the water cycle leads to increased salinity contrasts in
the ocean, both within and between basins. Regional projections for several areas important to large-scale
ocean circulation are presented, including the export of atmospheric moisture across the tropical Americas
from Atlantic to Pacific Ocean, the freshwater gain of high-latitude deep water formation sites, and the basin
averaged evaporation minus precipitation with implications for interbasin mass transports.
1. Introduction
The movement of water between oceanic, atmospheric, and land reservoirs is a complex component of
Earth’s climate collectively referred to as the water cycle. These flows move not only water but also heat
throughout the climate system. While it is the small
fraction of evaporation and precipitation that occurs
over land that directly impacts human populations and
industries, observational estimates indicate that 85% of
the global evaporation and 77% of precipitation occur
over the ocean (Durack 2015; Schanze et al. 2010;
Trenberth et al. 2007). This net evaporation over the
oceans is crucial for driving the relatively small amount
of precipitation over land (Gimeno et al. 2010, 2011; van
der Ent and Savenije 2013), and as such a full understanding of the water cycle requires a global perspective. Patterns of evaporation minus precipitation (E 2 P)
also have cascading impacts on oceanic circulations via
Corresponding author address: Samuel J. Levang, Dept. of Earth,
Atmospheric, and Planetary Sciences, 54-1611, Massachusetts Institute of Technology, 77 Massachusetts Ave, Cambridge, MA 02139.
E-mail: slevang@mit.edu
DOI: 10.1175/JCLI-D-15-0143.1
Ó 2015 American Meteorological Society
buoyancy and mass forcing, and on atmospheric circulation through latent heat release and cloud feedbacks
(Allan 2012).
The starting point for relating changes in global temperature to the water cycle is the increased mobility of
water molecules at higher temperatures. According to
the Clausius–Clapeyron (CC) relation for the saturation
vapor pressure es of water (Held and Soden 2006),
d ln(es )
L
5 y2 ’ 0:07 K21 .
dTs
RTs
Thus water vapor content in the atmosphere increases
by 7% K21 of warming at constant relative humidity for
Earth’s current mean temperature [Ly is the latent heat
of vaporization (2.5 3 106 J kg21), R is the gas constant
(461 J K21 kg21), and Ts is the near-surface air temperature]. Current global climate models (GCMs) obey CC
scaling quite closely for the increase in global atmospheric
moisture W (Held and Soden 2006; Allan et al. 2014).
The water cycle is made up of all transports and fluxes
of freshwater, and therefore ‘‘intensification’’ of the
water cycle is generally used to describe, in addition to
globally increased W, an increase in total precipitation
and evaporation, amplification of the existing E 2 P
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JOURNAL OF CLIMATE
pattern, and enhanced atmospheric vapor transport Q.
There is robust theoretical (Chou and Neelin 2004;
Muller and O’Gorman 2011), model (Held and Soden
2006; Chou et al. 2009), and even emerging observational (Huntington 2006; Durack et al. 2012) evidence
that the water cycle will intensify in the aforementioned
ways in a warming climate. However, the strength of the
water cycle is governed by more complex processes than
CC scaling. While there is a positive correlation between
global temperature and precipitation, the sensitivity is
thought to be less than CC (Stephens and Ellis 2008) due
to energetic constraints governing tropospheric radiative cooling (O’Gorman et al. 2012; Allan et al. 2014).
Analyses from phase 2 of the Coupled Model Intercomparison Project (CMIP) (CMIP2; Allen and Ingram
2002), phase 3 of CMIP (CMIP3; Held and Soden 2006),
and phase 5 of CMIP (CMIP5; Allan et al. 2014; Hegerl
et al. 2015) have shown the increase in global precipitation to be 2%–3% K21 with significant model
scatter.
Detecting changes in the water cycle via observations
of E 2 P has proved challenging, as the small net flux is a
difference between large and uncertain quantities
(Huntington 2006; Schmitt 2008; Schanze et al. 2010;
Josey et al. 2013). It has been suggested that a more
reliable indicator of the water cycle is sea surface salinity
(SSS), which is very sensitive to changes in the surface
fluxes. Evidence from 50 years of SSS observations
combined with CMIP3 scaling by Durack et al. (2012)
suggests that the sensitivity of SSS could be much higher
than CC and also more than predicted by climate
models, as much as 16% K21. Ocean mixing and advection work to smooth out the spatial and temporal
variability of surface water fluxes, and hence SSS is also
an effective integrator of long-term changes in the water
cycle (Durack et al. 2012). Analysis of observational
data has already revealed robust trends in global salinity
in line with projections of anthropogenic climate change
(Boyer et al. 2005; Stott et al. 2008; Durack and Wijffels
2010; Durack et al. 2012; Terray et al. 2012; Pierce et al.
2012). Preexisting interbasin salinity contrasts are
driven by net imbalances in E 2 P (Stommel 1980; Zhou
et al. 2000), and so intensification of the water cycle
tends to increase these basin contrasts.
Given the complexity of the water cycle, the inherent
coupling between oceanic and atmospheric processes,
and the sparseness of oceanic observations, GCMs are a
useful tool for predictive study of the water cycle.
Current-generation GCMs used in CMIP5 (Taylor et al.
2012) are generally effective at replicating global
patterns of E 2 P and atmospheric and oceanic transports of water. Liepert and Lo (2013) found that, with
several exceptions, CMIP5 models maintain a balanced
VOLUME 28
atmospheric water cycle without long-term drift in global
water content.
Several caveats should be noted regarding the representation of the water cycle in GCMs. First, although
model resolution has improved in recent generations,
current grid sizes still require significant topographical
smoothing. Therefore, steep mountain ranges such as
the Andes are represented with reduced peak heights,
and narrow oceanic throughflows like the Bering Strait
and Strait of Gibraltar have greatly simplified geometry.
A modeling study by Schmittner et al. (2011) demonstrated that alterations to global topography can substantially change water vapor fluxes, ocean circulation,
and the climate system as a whole. The highest South
American peak in models presented here ranges from
3916 to 4591 m, while the Andes reach an actual height
of 6962 m. This discrepancy may result in excess moisture
transport across ranges such as the Andes, as discussed by
Richter and Xie (2010). To combat these problems,
modelers sometimes enhance orography at the expense of
increasing the mean height of continents, resulting in a
dry bias of atmospheric vapor content over land (Gaffen
et al. 1997). In general, however, GCMs display a slightly
overactive water cycle, with greater total precipitation
and greater rainfall frequency than observations
(Stephens et al. 2010; Tian et al. 2013; Demory et al. 2014).
Nearly all current GCMs have well-documented issues with simulating tropical precipitation dynamics,
particularly in the Pacific Ocean (Lin 2007; Pincus et al.
2008; Collins et al. 2011; Brown et al. 2013; Flato et al.
2013; Li and Xie 2014). A set of connected biases combine to produce an excessive cold tongue in the eastern
Pacific, overly strong low-level trade winds, and an unrealistic double ITCZ flanking the equatorial Pacific
(Lin 2007; Hwang and Frierson 2013).
While all GCMs solve water conservation equations
and should ultimately produce an internally balanced
water cycle, some CMIP3 and CMIP5 models display
constant atmospheric moisture content despite having
globally unbalanced evaporation and precipitation,
suggesting an unphysical ‘‘ghost’’ source/sink of atmospheric moisture (Liepert and Previdi 2012; Liepert and
Lo 2013). The ocean component of GCMs can also
produce unrealistic freshwater balances, evidenced by a
drift in global mean salinity (Durack et al. 2014).
Therefore, water cycle consistency checks are important
when analyzing GCM data.
2. Theory and methods
a. Atmosphere
In a steady-state atmosphere, the vertically integrated
water vapor budget can be written as a simple balance
15 AUGUST 2015
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LEVANG AND SCHMITT
between horizontal transport divergence and the surface
flux E 2 P (following Peixóto and Oort 1983; Zaucker
and Broecker 1992; Cohen et al. 2000; Sohn et al. 2004;
Richter and Xie 2010):
$ Q 5 E 2 P,
where Q is the vertically integrated horizontal water
vapor transport, given by
Q5
1
g
ð top
qv dp,
surface
where q is the specific humidity, v is the horizontal wind
vector, p is the pressure, and g is the gravitational
constant.
Over any area, water gained or lost by E 2 P must
equal the transport through the boundaries:
þ
(Q n) dl 5 E 2 P dA,
∯
where dl is a line segment along the boundary and dA is
an element of surface area within the boundary. Note
that Q can be separated into mean and eddy components
via Reynolds decomposition:
Q 5 Q 1 Q0 .
For the zonal transport, particularly in the tropics
where winds are steadiest, Q0 is small (Wang et al. 2013).
However, in the meridional flux Q0 and Q are of similar
magnitude. To fully capture the meridional transport
then, the data used for this calculation must be sampled
at higher resolution than typical time scales of synoptic
weather systems. In this study, 6-hourly model outputs
have been used.
b. Ocean
In the ocean, any imbalance in the surface fluxes
(evaporation, precipitation, runoff, and ice flow)
induces a barotropic flow known as the Goldsbrough
circulation (Goldsbrough 1933; Huang and Schmitt
1993; Huang 2005). These circulations sometimes oppose the wind-driven gyres, such as in the North Atlantic, where net evaporation drives a northward
interior flow and a southward western boundary current
(Huang and Schmitt 1993). Applying steady-state mass
conservation [following Wijffels et al. (1992), Stammer
et al. (2003), Talley (2008), and Schanze et al. (2010)],
$ rv 5 E 2 P 2 R ,
and integrating through successive zonal sections, we
can obtain the meridional transport in each basin
dictated by the surface fluxes to within an integration
constant Tb:
ðð
ðð
E 2 P 2 R dx dy 5
rv dx dz 1 Tb ,
where Tb is the interbasin transport, applied at the
meridional boundaries of each basin; R is runoff into
the ocean; r is the ocean density; v is the meridional
ocean velocity; and x, y, and z are the zonal, meridional, and vertical coordinates respectively. As discussed
by Wijffels et al. (1992), the Bering Strait (BS) is a
particularly useful point of reference, as flows within
the Pacific and Atlantic are zonally bounded from this
point until they reach the Pacific Indonesian Throughflow
(PIT) and Southern Ocean. For ease of comparison to
previous observational estimates (Wijffels et al. 1992;
Schanze et al. 2010), the BS transport has been used along
with the net E 2 P 2 R over the Arctic to impose integration constants for the Atlantic and Pacific, and the
effect of the PIT has been neglected.
In steady state, the global net precipitation over land
exactly equals the runoff into the ocean. Therefore, a
closed budget can be derived for the ocean climatology
using only E 2 P data, but the precise location of runoff
sources will not be correctly represented, shifting the
meridional transport curve (Dai and Trenberth 2002). A
calculation that utilizes river and ice fluxes gives the
most accurate representation of ocean freshwater
transports. In the meridional integrations presented
here the water flux into ocean (wfo) model output has
been used, which combines all freshwater fluxes into a
single field on the native ocean grid.
c. Latent heat transport
Because of water’s high latent heat of vaporization, the
water cycle carries a significant amount of the poleward
heat flux. The latent heat flux is usually calculated from
the atmospheric moisture transport (Huang 2005):
HL 5 Ly Q,
here Q is a mass flux (in kg s21) of freshwater. The resulting heat flux associated with the meridional vapor
transport is as large as 2.5 PW near 458S or about half of the
total energy transport there (Trenberth and Caron 2001;
Huang 2005; also see Fig. 6). While this flux is often attributed to the atmospheric moisture transport, it is really a
product of the coupled ocean–atmosphere water cycle, as
the ocean must carry the compensating return flow of
liquid water. Therefore, while the ocean mass transports
induced by the surface flux E 2 P 2 R are an order of
magnitude smaller than the wind-driven circulation, they
are of great interest to the global energy budget.
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TABLE 1. Models utilized for the various projections and figures in this paper, listed by CMIP5 variable name (tas: near-surface air
temperature, hus: specific humidity, so: sea water salinity, evspsbl: evaporation, pr: precipitation, ua: zonal wind, and va: meridional wind).
Further model and variable specifications can be found in Flato et al. (2013). (Expansions of acronyms are available online at http://www.
ametsoc.org/PubsAcronymList.)
Institute
BCC
BCC
Beijing Normal University (BNU)
CCCma
CMCC
CMCC
CMCC
Centre National de Recherches Météorologiques
(CNRM)–Centre Européen de Recherche
et de Formation Avancée en Calcul
Scientifique (CERFACS)
CSIRO and BoM
CSIRO and BoM
CSIRO–Queensland Climate Change Centre of
Excellence (QCCCE)
First Institute of Oceanography (FIO)
Irish Center for High-End Computing (ICHEC)
Institute of Numerical Mathematics (INM)
IPSL
IPSL
IPSL
LASG and Center for Earth System Science (CESS)
JAMSTEC, Atmosphere and Ocean Research
Institute (AORI, The University of Tokyo),
and National Institute for Environmental
Studies (NIES)
JAMSTEC, AORI, and NIES
JAMSTEC, AORI, and NIES
Met Office Hadley Centre (MOHC)
MOHC
Max Planck Institute for Meteorology (MPI-M)
MPI-M
Meteorological Research Institute (MRI)
NASA GISS
NASA GISS
NASA GISS
NASA GISS
NCAR
Norwegian Climate Centre (NCC)
NCC
NOAA/GFDL
NOAA/GFDL
NOAA/GFDL
National Science Foundation (NSF)–DOE–NCAR
NSF–DOE–NCAR
tas and
hus
(Fig. 1)
so
(Fig. 3)
BCC_CSM1.1
BCC_CSM1.1(m)
BNU-ESM
CanESM2
CMCC-CESM
CMCC-CM
CMCC-CMS
CNRM-CM5
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
ACCESS1.0
ACCESS1.3
CSIRO Mk3.6.0
x
x
x
FIO-ESM
EC-EARTH
INM-CM4
IPSL-CM5A-LR
IPSL-CM5A-MR
IPSL-CM5B-LR
FGOALS-g2
MIROC-ESM
x
x
x
x
x
x
x
x
MIROC-ESM-CHEM
MIROC5
HadGEM2-CC
HadGEM2-ES
MPI-ESM-LR
MPI-ESM-MR
MRI-CGCM3
GISS-E2-H
GISS-E2-H-CC
GISS-E2-R
GISS-E2-R-CC
CCSM4
NorESM1-M
NorESM1-ME
GFDL CM3
GFDL-ESM2G
GFDL-ESM2M
CESM1(BGC)
CESM1(CAM5)
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
Model
d. Models and calculation specifics
For the projections presented here, all CMIP5 models
available on a local server (Table 1) were used for global
projections of temperature, vertically integrated water
vapor, evaporation, precipitation, and ocean salinity
evspsbl
and pr
(Fig. 2)
Q (ua, va, hus),
evspsbl, and pr
(Fig. 4) and
wfo (Figs. 5–7)
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
using monthly means (Figs. 1–4). A total of six models
have been used for the atmospheric and oceanic budget
and transport calculations (Figs. 5–7). The limited
number of models utilized in the transport calculations
is due to the large volume of data required when processing 6-hourly fields, which are required to adequately
15 AUGUST 2015
LEVANG AND SCHMITT
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FIG. 1. The 100-yr projection [from historical (1990–2000) to RCP8.5 (2090–2100)] for (a) change in surface air
temperature Ts and (b) percent increase in vertically integrated water vapor content W. Global average temperature change is 4.08C. The warming closely matches the increase in water vapor content, with a pattern correlation
of 0.86. The projections are an average of the 38 models listed in Table 1.
capture the eddy transports in the atmosphere. Only
models found by Liepert and Lo (2013) to have realistically balanced water cycles were selected for the
budget and transport calculations. Models with a range
of horizontal grid sizes were intentionally chosen to
explore the effect of topographical smoothing on the
atmospheric transport. All multimodel means (MMM)
use a single ensemble run (r1i1p1) from each model.
Ten-year averages from the historical (1990–2000) and
RCP8.5 (2090–2100) scenarios were processed to quantify
centennial-scale changes in the water cycle under global
warming. All ‘‘projections’’ described in the results section are taken to be MMM changes between these two
time periods. Representative concentration pathway
8.5 (RCP8.5) is a high-emissions scenario intended to
produce an additional radiative forcing of 8.5 W m22 by
the year 2100, although the effective forcing varies between models and was found by Forster et al. (2013) to be
only 7.4 W m22 in the MMM. RCP8.5 produces a rise in
MMM surface air temperature of 4.08C for the 100 years
analyzed. It is used here to represent the potential for
changes in the water cycle but should not be interpreted
as a quantitatively accurate forecast for the coming century. It should be noted that the CC scaling associated
with this temperature rise would predict a 28% increase in
surface specific humidity.
Additionally, a water budget has been calculated from
state-of-the-art observationally based products from the
satellite era (1987–2014) as a comparison for the CMIP5
historical simulations. The atmospheric moisture transport Q was calculated from ERA-Interim (Dee et al.
2011) reanalysis output, and the estimate of surface flux
into the ocean E 2 P 2 R was compiled from GPCP
precipitation (Adler et al. 2003), OAFlux evaporation
(Yu and Weller 2007), and runoff estimates from Dai
and Trenberth (2002) [see Schanze et al. (2010) for selection]. This E 2 P 2 R estimate for the ocean gives
total P of 12.6 Sv, E of 13.1 Sv, and R of 1.25 Sv, with an
imbalance of only 0.13 Sv (1 Sv [ 106 m3 s21).
The Q calculation was performed using model outputs
at 6-hourly time steps for wind and humidity in order to
fully capture the eddy transport. The outputs were
processed in the native hybrid-sigma pressure coordinate, which conforms to topography at low altitudes,
and a trapezoidal scheme was used for the discrete
vertical integral. The horizontal output was interpolated
onto a standardized 18 3 18 grid for each model. ERAInterim offers a direct monthly output of Q.
Ocean freshwater fluxes for the meridional integration were taken from the combined E 2 P 2 R flux
into the ocean of each model and first integrated on the
native ocean grid, in order to avoid introducing interpolation errors at point source river mouths.
3. Results and projections
RCP8.5 produces a rise in global mean surface air
temperature of 4.08C and an increase in atmospheric
water content W of 32%, versus the predicted 28% for
CC scaling (Fig. 1). Model projections for E 2 P are
presented in Fig. 2. Total global evaporation (equal to
global precipitation) increases from 17.3 to 18.6 Sv in the
multimodel 100-yr projection, an 8% increase, or about
2% K21. Best observational estimates based on satellite
data from GPCP (Trenberth et al. 2007) estimate global
precipitation at 15.4 Sv. The net water transport from
ocean to land (E 2 P over land, equivalent to R) increases by 12% from 1.17 to 1.31 Sv. Total precipitation
increases by 10% over land and just 6% over the ocean.
A greater acceleration of the water cycle over land is
expected due to slow thermal adjustment of the oceans
(Dommenget 2009) and subsequently greater warming
over land in the 100-yr period analyzed here.
While evaporation and precipitation both increase
globally, planetary-scale patterns of E 2 P also show
coherent changes. There is high model agreement
(.90%) that ocean locations poleward of 458N and
458S will experience increased precipitation. The five
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FIG. 2. (a) The E 2 P (m yr21) for the historical period (1990–2000). The 100-yr projection for change in
(b) E 2 P, (c) E, and (d) P (m yr21), using RCP8.5 (2090–2100) and historical (1990–2000) simulations. Areas
with greater than 90% model agreement in the sign of the change are stippled. All panels are an average of the
28 models listed in Table 1.
subtropical ocean gyres display a decrease in precipitation albeit with weaker model agreement. There is
high agreement for an increase in evaporation over almost all of the oceans except the Southern Ocean, which
also shows reduced warming relative to the global mean.
Projections for land areas show a modest increase in
net precipitation but with poor model agreement. The
model average net precipitation increases significantly
in the tropical rain belt of the Pacific and Indian Oceans
but with large uncertainty among models. Integrated
over each basin, the Atlantic becomes 0.20 Sv more
evaporative, the Pacific 0.16 Sv more precipitative, the
Indian Ocean 0.08 Sv more evaporative, and the Arctic
Ocean 0.05 Sv more precipitative.
SSS is strongly forced by the surface flux E 2 P 2 R,
and hence would be expected to display similar
changes. The meridional patterns of increased subtropical
evaporation and increased tropical and high-latitude
precipitation are evident in the SSS projection presented in Fig. 3, but are dwarfed by a large zonal contrast
between subtropical Atlantic and Pacific. With the exception of modest salinification in the southern subtropical Pacific, all other areas of the Pacific display
surface freshening. The weak signature in the northern
subtropical Pacific may partly be a function of weakened
trade winds and evaporation (La^ıné et al. 2014), as well as
advection of freshened water from the equatorial Pacific.
In the Atlantic, all subtropical regions show surface salinification, as high as 0.8. As the area of the low-latitude
Atlantic region defined here is less than half of its Pacific
counterpart, it is expected that the surface salinity forcing
would be felt more strongly in the Atlantic.
Following Durack et al. (2012), an MMM linear regression of zonal mean SSS change onto its modeled
FIG. 3. (a) Multimodel mean climatology for SSS. (b) Projection for 100-yr change in sea surface salinity, using
RCP8.5 (2090–2100) and historical (1990–2000) simulations. Salinities are given in practical salinity scale (PSS-78)
equivalent. Areas where more than 90% of models agree on the sign of the change are stippled. Both panels are an
average of the 31 models listed in Table 1.
15 AUGUST 2015
LEVANG AND SCHMITT
historical values gives a pattern amplification of 5.9% K21,
slightly lower than their values for CMIP3. The same
calculation for E 2 P gives a pattern amplification of
5.2% K21, in line with previous values (Durack et al.
2012). The SSS regression is particularly sensitive to the
use of individual model climatologies versus MMM climatology, as performing the regression on the MMM
smooths out covariances and gives only 3.1% K21.
While the projected E 2 P changes do not at first
glance display this zonal asymmetry, salinity is an effective spatial integrator of surface forcing as ocean
currents and mixing tend to smear out smaller-scale
variability. It is therefore useful to assume a basinwide
perspective of the atmospheric transports and their associated freshwater fluxes into and out of the oceans.
Figure 4 summarizes this perspective, with drainage
basins defined by topographical maxima. Low- and highlatitude basins are also divided at 308N and 308S. The
transport from tropical Atlantic to Pacific has been broken into three components: South America (308S–88N),
Central America (88–208N), and Mexico (208–308N).
The atmospheric water budget produced by the
CMIP5 historical simulations (Fig. 4b) shows many
similarities to the observational estimate (Fig. 4a). Both
capture the export of water vapor from low to high latitudes, zonal convergence in the Pacific, and divergence
in the Atlantic and Indian Ocean. The models and observations produce the opposite sign for the small net
flux across the Maritime Continent. The largest discrepancies seen in the fluxes are across South America,
with the CMIP5 models producing a much higher
transport estimate there. There is also a significant discrepancy in the GPCP/OAFlux and ERA-Interim budget for the low-latitude Pacific with a residual of 0.28 Sv,
making this region poorly constrained for comparison
with models. An imbalance of 0.27 Sv is also present in
the Southern Ocean region. There are small imbalances
between the CMIP5 divergences and surface fluxes as
well, but all are less than 0.1 Sv. These model imbalances
are likely related to the regridding scheme used here, or
the relatively short 10-yr integration which may allow
some reservoir storage.
The RCP8.5 projection for basin to basin transports
reveals the larger-scale pattern that explains the zonal
SSS trend in Fig. 3. Relative to the historical scenario,
RCP8.5 shows an increased westward transport by the
trade winds at low latitudes, and an increased poleward
transport from low to high latitudes driven primarily by
the synoptic eddy field. Although the zonal transport
across all low-latitude divides increases, the preexisting
differences in magnitude are amplified by such a pattern.
Therefore the net evaporation in the low latitudes gets
larger in the Atlantic and smaller in the Pacific.
6495
Additionally, there is an increase in net precipitation for
all extratropical regions.
Among the six models analyzed here, there is 100%
agreement for the sign of the change in net E 2 P 2 R
for all basins except the tropical Indian Ocean (5/6 in
agreement). There is similarly robust agreement for the
changes in atmospheric transport across drainage divides,
with many segments below displaying 100% agreement in
the sign of the change. The components with some model
disagreement are the northward export out of the tropical
Atlantic (4/6 in agreement), the southward export out of
the tropical Pacific (4/6 in agreement), the tropical Pacific
to tropical Indian Ocean (5/6 in agreement), and across
the Himalayas (3/6 in agreement).
The MMM of zonally averaged transports (Fig. 5)
shows an increase in atmospheric freshwater fluxes
everywhere. The intermodel spread is displayed for an
idea of the model uncertainty in moisture transports.
The change in the zonally integrated meridional transport (Fig. 6) is as large as 0.3 Sv near 108 and 408S, which
corresponds to nearly 0.7 PW of additional latent heat
transport. The meridional ocean transports separated
into basins (Fig. 7) show generally weaker southward
transport through the Atlantic and weaker northward
transport through the Pacific. All transport curves in
Figs. 5–7 match qualitatively with the observational
comparison.
4. Discussion
a. Zonal atmospheric transports
To first order, Q scales similarly to CC, which would
be the expectation if warming caused no dynamical response in the circulation (Held and Soden 2006). The
global mean zonal transport Qx increases 35%, and the
meridional transport Qy increases 30%, noting again
that the increase in W is 32%. A robust consequence of
this thermodynamic scaling is that regional patterns of
moisture convergence and divergence are amplified in a
warming climate even without circulation changes. The
preexisting zonal pattern in the subtropics is a useful
example. There is a large moisture transport across
Central and South America relative to the transport
across Africa and the Maritime Continent. If the increase in temperature and thus specific humidity is
uniform, then a constant fractional increase of each interbasin transport causes an increased volume convergence in the Pacific.
A slight increase in the strength of the trades over
South America causes the change in Q to be higher than
CC scaling. Conversely, a slight decrease in the trades
across Africa gives a lower scaling there. In the models
analyzed here, there is not a significant correlation
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FIG. 4. Summary of the atmospheric freshwater budget showing Q and E 2 P for (a) ERAInterim winds and humidity with GPCP/OAFlux fluxes for 1979–2014, and multimodel means
of the CMIP5 (b) historical (1990–2000) and (c) RCP8.5 (2090–2100) simulations. For the
observational dataset, E 2 P 2 R into the ocean has been used rather than E 2 P over combined land and ocean areas. The map has been divided (thick black lines) by topographical
drainage basins between the Atlantic, Pacific, and Indian Oceans, and into low- and highlatitude regions separated at 308N and 308S, corresponding to the approximate center of the
Hadley cells. Colored values are E 2 P integrated over each region, with red indicating net
evaporation and blue net precipitation (Sv). Filled contours are the global E 2 P (m yr21).
Small arrows indicate the vertically integrated atmospheric moisture flux, and thick arrows
show the fluxes across drainage divides (Sv). The CMIP5 transports are an average of the six
models listed in Table 1.
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LEVANG AND SCHMITT
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FIG. 5. Zonally averaged meridional and zonal atmospheric vapor transport, for ERAInterim (thin solid lines), historical (thick solid lines), and RCP8.5 (dashed lines) simulations.
The transports are expressed as a flux density per meter (kg m21 s21) of latitude. The zonal
transports are an order of magnitude larger than the meridional ones. The CMIP5 transports
are an average of the six models listed in Table 1.
between horizontal resolution and transport over the
Andes, although all display much greater transports
than the ERA-Interim value of 0.16 Sv. Other reanalysis
products show greater Andean transports closer to the
CMIP5 values (Richter and Xie 2010), particularly along
the northern segment, so it is unclear how topographical
smoothing affects this transport. Xu et al. (2005) showed
that, in a high-resolution model, transport across Central America was insensitive to the fine structure of the
orography, as intense jets around mountains or diffuse
flow over smooth topography produced nearly the same
transport.
There are a number of dynamical factors affecting
tropical transports in a warming climate, although
circulation changes are generally small compared to
thermal affects (Held and Soden 2006). A common
feature of GCM warming simulations is a decrease in the
Walker circulation (Vecchi et al. 2006; DiNezio et al.
2013; Ma and Xie 2013). This is sometimes explained
using the differential response of atmospheric water
vapor (7% K21) and precipitation (2%–3% K21) to
warming. Because the convective supply of water cannot
exceed the loss of water from precipitation, the balance
equations require a weakening of the upward mass flux
and therefore the overturning circulation (Held and
Soden 2006; Ma et al. 2012). However, observational
evidence for changes in the Walker circulation is inconclusive (Tokinaga et al. 2012).
FIG. 6. Zonally integrated meridional freshwater transport (Sv) in the atmosphere and the
ocean, for observations (thin solid lines), historical (thick solid lines), and RCP8.5 (dashed
lines) simulations. The atmospheric calculation uses the zonal integral of Q and the ocean curve
is produced from the E 2 P 2 R integral or wfo variable in CMIP5. The CMIP5 transports are
an average of the six models listed in Table 1.
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VOLUME 28
FIG. 7. The E 2 P 2 R derived freshwater transports in each basin, using the Bering Strait
flow as an integration constant for the Atlantic and Pacific. The observational curves are the
integrated E 2 P 2 R dataset, and the CMIP5 curves are an average of the wfo variable for the
six models listed in Table 1. The net Arctic freshwater gain is added to the southward Atlantic
flow at 658N: 0.13 Sv for the observations, 0.20 Sv for the historical simulation, and 0.29 Sv for
RCP8.5.
Regional patterns of SST increase may also play a role
(Knutson and Manabe 1995; Tokinaga et al. 2012). In
the tropics, a distinct El Niño–like warming mode is
produced by CMIP5 models (Cai et al. 2014; Fig. 1a).
Warm SST in the eastern tropical Pacific induces
anomalous low pressure there, and an increase in trade
wind strength across South America follows.
The change in Q across the Maritime Continent
(88%) is large in a relative sense, but remains small in
terms of volume flux because the net transport of the
seasonal monsoonal cycle is small. One argument regarding the monsoon circulation of the Indian Ocean
basin predicts an increase in summer monsoon intensity
due to the transient increase in land–sea temperature
contrast under warming (Turner and Annamalai 2012;
Lee and Wang 2014). Such an increase would bring additional moisture from the Pacific to the Indian Ocean
during summer, but monsoon projections have large
intermodel spread. The change in Q is also large (44%)
for the extratropical transport across North America
from the Pacific to the Atlantic, dominated by anomalous northeasterly transport over Alaska, but this is
primarily a function of amplified Arctic warming and
therefore larger CC scaling.
Based on these trends in interbasin transport, a general enhancement of the zonal pattern of E 2 P and SSS
is projected under anthropogenic warming. The Atlantic
becomes more evaporative and saline, while the Pacific
becomes more precipitative and fresher. These trends
will lead to halosteric effects on sea level, diminishing
sea level rise in the Atlantic while enhancing it in the
Pacific (Durack et al. 2014). Despite an increase in net
evaporation over the Indian Ocean, a slight freshening is
projected for all but the most western portions of the
basin. This is consistent with the substantial advection of
Pacific water westward through the PIT, water that is
projected to freshen because of increased precipitation
in the western Pacific.
Increased moisture export from the Atlantic to the
Pacific and subsequent salinification of Atlantic surface
waters has previously been discussed as a positive
feedback on North Atlantic Deep Water (NADW)
formation and the meridional overturning circulation
(Broecker et al. 1990; Zaucker et al. 1994; Lohmann
2003; Richter and Xie 2010). Increased freshening and
warming of the northern Atlantic and Arctic surface
water tends to increase static stability of the water column and decrease deep convection (Aagaard and
Carmack 1989; Hasumi 2002; Hu and Meehl 2005), and
CMIP5 models generally do show a reduction in the
Atlantic overturning. However, these simulations also
confirm that warming increases surface salinities in the
subtropical Atlantic, which could counteract the positive buoyancy forcing from increased freshwater input at
higher latitudes.
b. Meridional atmospheric transports
The projected increase in meridional transport across
308N and 308S is somewhat less than CC scaling, and also
shows large zonal variation. The synoptic eddy field
plays a much greater role in determining meridional
fluxes compared to zonal fluxes, and so the meridional
transports may be more subject to dynamical changes
and less directly tied to CC scaling. Poleward transport
out of the North Pacific increases by 31%, while transport out of the North Atlantic shows nearly no change,
15 AUGUST 2015
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LEVANG AND SCHMITT
as increases in W are offset by a weakening of the prevailing southwesterly transport there. In the Southern
Hemisphere, the poleward transports out of the Atlantic
and Indian Oceans are near CC scaling. Transport out of
the South Pacific, however, shows almost no change.
Here CC is counteracted by an equatorward anomaly
along South America. Transport across the Himalayas is
negligible for both scenarios.
c. Ocean throughflows and transports
Bering Strait (BS) transport is understood to be a
function of the sea surface height (SSH) gradient from
Pacific to Atlantic, which some have suggested is
maintained by atmospheric moisture export over Central and South America and associated E 2 P fluxes
(Latif 2001; Schmitt 2008). Despite projections for increasing disparity between Pacific and Atlantic E 2 P,
the six models analyzed here show a decreasing BS flow,
from 0.80 to 0.68 Sv in the MMM, compared to observational estimates of 0.79 Sv (Schanze et al. 2010). While
the Atlantic becomes more evaporative and would be
expected to decrease in mean height from mass balance
and halosteric effects, the Arctic experiences significant
freshwater gain, which increases SSH there. Wind stress
has been used to explain seasonal variability in the Bering Strait transport (Coachman and Aagaard 1988), as
the mean wind field is southward and opposes the
pressure-driven ocean transport. Wind stress does not
appear to explain the projections for Bering Strait
transport here, as the change in mean wind fields is actually for weaker southerly circulation through the
strait. A probable explanation for decreased BS flow is
the thermosteric rise associated with amplified warming
in the Arctic, reducing the more local Pacific to Arctic
pressure gradient.
Basin meridional transports for the historical simulations closely resemble those of observations
(Baumgartner and Reichel 1975; Wijffels et al. 1992; Dai
and Trenberth 2003, Schanze et al. 2010). The PIT
transport has been neglected in Fig. 7 for simplicity,
which induces a counterclockwise circulation around
Australia [see Wijffels et al. (1992) and Dai and
Trenberth (2003) for further discussion] and would significantly displace the Indian Ocean curve toward
southward transport and the Pacific curve toward
northward transport south of the PIT. The observational
estimate presented here shows stronger northward
transport in the Pacific and weaker southward transport
in the Atlantic from approximately 108N and all points
southward relative to CMIP5. Because of the slight decrease in BS flow and increased precipitation in the
tropical Pacific, the projected meridional flow in the
Pacific becomes southward around the equator, unlike
in the historical simulation, which produces northward
flow throughout the Pacific.
5. Summary
This study has examined the estimated changes in the
water cycle of relevance to ocean circulation from a
multimodel ensemble of CMIP5 models for a high CO2
emissions scenario for the twenty-first century. Drainage
basins are defined and net water fluxes calculated between them. There are large changes in interbasin transports and strengthening of the patterns of midlatitude
evaporation and high- and low-latitude precipitation.
Substantial changes in ocean salinity are projected, with
salty areas getting saltier and fresh areas fresher overall.
Such changes have consequences for regional sea level
rise and the strength of the meridional overturning circulation, and indicate the importance of sustained global
salinity observations over the next century. However, the
spatial response of SSS and E 2 P to warming in CMIP5 is
far from identical. These results suggest that if the salinity
field is to be used as a global ‘‘rain gauge’’ for Earth’s
water cycle, the effects of ocean mixing and advection and
large-scale interbasin patterns of moisture convergence
must be considered in addition to the predicted amplification of the local surface forcing.
Acknowledgments. This research was supported by
NASA Grant NNX12AF59GS03. The authors thank all
of the modeling groups involved in CMIP5, the World
Climate Research Programme’s Working Group on
Coupled Modeling for coordinating the experiment, and
those involved in the WHOI CMIP5 Community Storage
Server supported by the Ocean Climate Change Institute.
In addition, the creators of the observational datasets
used here including the WHOI OAFlux Project, NASA
GEWEX Global Precipitation Climatology Project, and
the ECMWF interim reanalysis (ERA-Interim) are acknowledged. Finally, the authors thank two anonymous
reviewers for their exceptionally helpful comments.
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