Toward a Unified View of the American Monsoon Systems

C. VERA,a W. HIGGINS,b J. AMADOR,c T. AMBRIZZI,d R. GARREAUD,e D. GOCHIS,f D. GUTZLER,g

D. LETTENMAIER,h J. MARENGO,i C. R. MECHOSO,j J. NOGUES-PAEGLE,k P. L. SILVA DIAS,d AND

C. ZHANGl

a CIMA/University of Buenos Aires, Buenos Aires, Argentina
b Climate Prediction Center, NOAA/NWS/NCEP, Camp Springs, Maryland

c University of Costa Rica, San Jose, Costa Rica
d University of Sao Paulo, Sao Paulo, Brazil

e Universidad de Chile, Santiago, Chile
f National Center for Atmospheric Research/RAP, Boulder, Colorado

g University of New Mexico, Albuquerque, New Mexico
h University of Washington, Seattle, Washington

i CPTEC/INPE, Cachoeira Paulista, Brazil
j University of California, Los Angeles, Los Angeles, California

k University of Utah, Salt Lake City, Utah
l RSMAS, University of Miami, Miami, Florida

(Manuscript received 31 August 2004, in final form 14 February 2005)

ABSTRACT

An important goal of the Climate Variability and Predictability (CLIVAR) research on the American
monsoon systems is to determine the sources and limits of predictability of warm season precipitation, with
emphasis on weekly to interannual time scales. This paper reviews recent progress in the understanding of
the American monsoon systems and identifies some of the future challenges that remain to improve warm
season climate prediction. Much of the recent progress is derived from complementary international pro-
grams in North and South America, namely, the North American Monsoon Experiment (NAME) and the
Monsoon Experiment South America (MESA), with the following common objectives: 1) to understand the
key components of the American monsoon systems and their variability, 2) to determine the role of these
systems in the global water cycle, 3) to improve observational datasets, and 4) to improve simulation and
monthly-to-seasonal prediction of the monsoons and regional water resources. Among the recent obser-
vational advances highlighted in this paper are new insights into moisture transport processes, description
of the structure and variability of the South American low-level jet, and resolution of the diurnal cycle of
precipitation in the core monsoon regions. NAME and MESA are also driving major efforts in model
development and hydrologic applications. Incorporated into the postfield phases of these projects are
assessments of atmosphere–land surface interactions and model-based climate predictability experiments.
As CLIVAR research on American monsoon systems evolves, a unified view of the climatic processes
modulating continental warm season precipitation is beginning to emerge.

1. Introduction

Monsoon circulation systems, which develop over
low-latitude continental regions in response to seasonal
changes in the thermal contrast between the continent
and adjacent oceanic regions, are a major component of

continental warm season precipitation regimes. Both
North and South America are characterized by such
systems [hereafter referred to as the North American
Monsoon System (NAMS) and the South American
Monsoon System (SAMS), respectively]. The NAMS
and SAMS provide a useful framework for describing
and diagnosing warm season climate controls, and the
nature and causes of year-to-year variability. A number
of studies during the past decade have revealed the
major elements of these systems, including their con-
text within the annual cycle, and some aspects of their
variability.

Corresponding author address: Dr. Carolina Vera, CIMA, Pab.
II, 2do piso, Ciudad Universitaria, (1428) Buenos Aires, Argen-
tina.
E-mail: carolina@cima.fcen.uba.ar

15 OCTOBER 2006 V E R A E T A L . 4977

© 2006 American Meteorological Society

JCLI3896



Due in large part to the success of the World Climate
Research Programme/Climate Variability and Predict-
ability/Variability of the American Monsoon Systems
(WCRP/CLIVAR/VAMOS) program, a unifying view
of the NAMS and SAMS is beginning to emerge. In
particular, CLIVAR/VAMOS has implemented
complementary international programs in North and
South America, namely the North American Monsoon
Experiment (NAME) and the Monsoon Experiment
South America (MESA). These programs have com-
mon objectives: 1) to understand the key components
of the American monsoon systems and their variability,
2) to determine the role of these systems in the global
water cycle, 3) to improve observational datasets, and
4) to improve simulation and monthly-to-seasonal pre-
diction of the monsoon and of regional water resources.

Several recent publications summarize the accom-
plishments of the scientific community toward achiev-
ing a better understanding of the monsoon systems of
the Americas. These include a review of the South
American Monsoon System (Nogués-Paegle et al.
2002) and a review of the North American Monsoon
System (Higgins et al. 2003). The intention of this paper
is to discuss recent advances in our understanding of
the NAMS and SAMS in an integrated framework,
highlighting recent papers that illustrate progress made
during the initial phase of CLIVAR. More complete
historical reference lists are found in the overview pa-
pers cited above.

The basic features of the NAMS and SAMS and their
variability within the context of the land surface–atmo-
sphere–ocean annual cycle are discussed in sections 2
and 3, respectively. The role of land surface memory in
the variability and predictability of the NAMS and
SAMS is considered in section 4. Hydrologic character-
istics of the monsoon systems are described in section 5.
Outstanding science questions associated with gaps in
our understanding are reviewed in section 6.

2. Basic features

Both the NAMS and the SAMS exhibit many of the
features of their Asian counterpart, including large-
scale land–sea temperature contrast, a large-scale ther-
mally direct circulation with a continental rising branch
and an oceanic sinking branch, land–atmosphere inter-
actions associated with elevated terrain and land sur-
face conditions, surface low pressure and an upper level
anticyclone, intense low-level inflow of moisture to the
continent, and associated seasonal changes in precipi-
tation (both increases and decreases). Moreover, the
poleward extension of the summer convection in the
tropical Americas seems to be associated with similar

mechanisms (such as the ventilation processes) to those
acting in the Asian monsoon system (Chou and Neelin
2003).

Both the NAMS and the SAMS receive more than
50% of total annual precipitation during the respective
summer monsoons, though the SAMS precipitation
amounts are considerably greater (Figueroa and Nobre
1990; Higgins et al. 1997). The SAMS exhibits some-
what distinct characteristics compared to the other
monsoon systems, given that most of South America is
situated in the Tropics, and seasonal temperature dif-
ferences are less pronounced than in subtropical mon-
soon regimes.

A clear annual cycle characterizes the convection
over the tropical Americas that exhibits a seasonal
regularity and degree of symmetry with respect to the
equator (Horel et al. 1989). The NAMS and SAMS can
be interpreted as the two extremes of the same cycle
(Fig. 1) and their corresponding life cycle can be de-
scribed using terms that have been traditionally re-
served for the Asian summer monsoon system, namely,
onset, mature, and decay phases.

Seasonal evolution of the convection

During May–June (the onset phase of the North
American monsoon), heavy rains spread northward
along the western slopes of the Sierra Madre Occiden-
tal (SMO; Fig. 2; Douglas et al. 1993; Stensrud et al.
1995; Adams and Comrie 1997). Precipitation increases
over northwestern Mexico coincide with increased ver-
tical transport of moisture by convection (Douglas et al.
1993) and southerly winds flowing along the Gulf of
California. During this season, weather in the NAMS
region changes abruptly from relatively hot, dry condi-
tions to cool, rainy ones (Mock 1996; Adams and Com-
rie 1997).

Increases in precipitation over the southwestern
United States occur abruptly around the beginning of
July (e.g., Mock 1996; Higgins et al. 1997) and coincide
with the development of a pronounced anticyclone at
the jet stream level (e.g., Okabe 1995), the develop-
ment of a thermally induced trough in the desert South-
west (Rowson and Colucci 1992), northward displace-
ments of the Pacific and Bermuda highs (Carleton
1987), the formation of southerly low-level jets over the
Gulf of California (Douglas 1995), and the formation of
the Arizona monsoon boundary (Adang and Gall
1989). The onset of the summer monsoon rains over
southwestern North America has been linked to an in-
crease of rainfall along the East Coast of the United
States (e.g., Higgins et al. 1997) and to a decrease of
rainfall over the Great Plains of the United States (e.g.,
Douglas et al. 1993; Mock 1996; Higgins et al. 1997).

4978 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19



Moisture transport onto the North American continent
is accomplished by boundary layer flow from the south
(Gulf of California), and in the middle troposphere via
southeasterly flow from the Gulf of Mexico (Schmitz
and Mullen 1996; Fig. 3a).

During July–August–early September, the NAMS is
fully developed. The heaviest precipitation is west of
the SMO and near the Bay of Campeche (Fig. 4b). The
northern edge of the monsoon extends into Arizona
and New Mexico (e.g., Douglas et al. 1993), but the
rainfall is much lighter and more directly influenced by
midlatitude effects. Precipitation also occurs in Central
America and northwestern South America, with a rela-
tive maximum of precipitation in southeast (SE) sub-
tropical South America (Fig. 4a). The diabatic heating
released by the NAMS in combination with the oro-
graphic forcing induces descent over the eastern North
Pacific while it promotes the development of the North
Atlantic subtropical high (Rodwell and Hoskins 2001).

Surges of maritime tropical air move northward
along the Gulf of California and are linked to bursts
and breaks of the monsoon rains over the deserts of
Arizona and California (Stensrud et al. 1995). Gulf
surges are triggered by a variety of synoptic-scale and
mesoscale disturbances, including tropical easterly
waves, tropical cyclones, mesoscale convective systems,
and upper-level inverted troughs (e.g., Stensrud et al.
1995; Higgins et al. 2004). Equatorward of the tropic of
Cancer the monsoon exhibits a double-peak structure
in precipitation and diurnal temperature range. From

south-central Mexico into Central America, this mid-
summer dry spell or “canicula” is sufficiently regular as
to appear in climatological averages (Magaña et al.
1999). Recent evidence indicates that the trade winds,
evaporation, and precipitation patterns over the warm
pool region to the southwest of Mexico modulate sea
surface temperatures (SSTs) in a manner consistent
with the double-peak structure in precipitation (Ma-
gaña et al. 1999).

During late September–October the decay phase of
the NAMS occurs. The ridge over the western United
States weakens as the monsoon high retreats southward
and precipitation diminishes. Simultaneously the mid-
latitude westerly regime shifts equatorward, and pre-
cipitation events are more frequently associated with
synoptic-scale frontal systems rather than with localized
convective instability.

By September, the convection migrates from Central
America into South America, and the onset of the wet
season over South America starts first in the equatorial
Amazon and then spreads quickly to the east and
southeast (Fig. 2b). The onset across the Amazon basin
lasts about one month (e.g., Kousky 1988; Horel et al.
1989; Marengo et al. 2001; Liebmann and Marengo
2001) and is followed by abundant rainfall. The onset
(demise) of the Amazon rainy season is preceded by an
increase in the frequency of northerly (southerly) cross-
equatorial flow over South America (Marengo et al.
2001; Wang and Fu 2002). Changes in the moistening of
the planetary boundary layer and changes in the tem-

FIG. 1. Mean annual cycle of precipitation over several major monsoon areas [NAMS
(20°–37°N, 248°–257°E); SAMS 40°–60°W); India (6°–37°N, 68°–98°E); Sahel (10°–20°N, 15°–
15°W)]. For comparison, one nonmonsoon region with a large annual cycle is also shown
[Pacific Northwest (PNW: 42°–50°N, 112°–124°W)].

15 OCTOBER 2006 V E R A E T A L . 4979

Fig 1 live 4/C



perature at the top of the boundary layer seem to ex-
plain the seasonal changes of convection in tropical
South America (Fu et al. 1999; Marengo et al. 2001;
Liebmann and Marengo 2001).

The onset of the wet season in central and southeast-
ern Brazil typically occurs between the end of Septem-
ber and early October and may sometimes occupy only
a single 5-day period (Sugahara 1991). Intraseasonal
oscillations may promote rapid onset over central Bra-

zil (Vera and Nobre 1999). By late November, deep
convection covers most of central South America from
the equator to 20°S, but is absent over the eastern
Amazon basin and northeast Brazil. Throughout this
period, deep convection associated with the intertropi-
cal convergence zone (ITCZ) is confined to the central
Atlantic between 5° and 8°N (Zhou and Lau 1998).

During late November through late February (the
mature phase of the SAMS), the main convective ac-
tivity is centered over central Brazil and linked with a
southeastward band of cloudiness and precipitation ex-
tending from southern Amazonia toward southeastern
Brazil and the surrounding Atlantic Ocean (Fig. 4b).
That convection band, known as the South Atlantic
convergence zone (SACZ), is a distinctive feature of
the SAMS (Kodama 1992). Also, the heavy rainfall
zone extends over the Altiplano Plateau and the south-
ernmost Brazilian highland.

The upper-level circulation during the South Ameri-
can summer includes well-defined regional circulation

FIG. 2. Mean calendar date of onset for (a) the North American
monsoon (from Higgins et al. 1999), and (b) the South American
Monsoon (courtesy of V. Kousky).

FIG. 3. (a) Schematic vertical (longitude–pressure) cross section
through the NAMS at 27.5°N. Topography data were used to
establish the horizontal scale and NCEP–National Center for At-
mospheric Research (NCAR) reanalysis wind and divergence
fields were used to establish the vertical circulations (from Hig-
gins and NAME Science Working Group 2003). (b) Schematic
vertical section across South America displaying the major large-
scale elements affecting the SAMS (from CLIVAR Web site on-
line at http://www.clivar.org).

4980 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19

Fig 2 3 live 4/C



features, including a large anticyclonic circulation (the
“Bolivian High”) centered near 15°S, 65°W and an up-
per-level trough near the coast of northeast Brazil (not
shown). At low levels, the “Chaco Low” is the most
conspicuous feature downslope of the Andes and can
be considered together with the Bolivian High as the
regional response of the tropospheric circulation to the
strong convective heating over the Amazon–central
Brazil. The Andes effect reinforces the strength of the
Chaco Low through the barrier role of the mountains
(Gandu and Silva Dias 1998, and references therein). A
continental-scale gyre transports moisture westward
from the tropical Atlantic Ocean to the Amazon basin,

and then southward toward the extratropics of South
America (Fig. 4a). The diabatic heating released in the
SAMS region seems to promote that gyre, and the
maintenance of the South Atlantic subtropical high
during austral summer (Rodwell and Hoskins 2001). A
regional intensification of this gyre circulation to the
east of the Andes Mountains is due to the South Ameri-
can low-level jet (SALLJ), with strongest winds in Bo-
livia near Santa Cruz (18°S, 63°W). The SALLJ trans-
ports considerable moisture between the Amazon and
the La Plata basins and is present throughout the year
(e.g., Berbery and Barros, 2002); it can be explained
using simple, adiabatic models in which orography pro-
vides dynamical, rather than thermodynamical modifi-
cation of the zonally averaged circulation (Byerle and
Paegle 2002; Campetella and Vera 2002). During the
warm season, thermodynamic processes associated with
precipitation either over the SACZ region or southeast-
ern South America (SESA) modulate the low-level
flow in tropical regions (Berbery and Collini 2000).

Between March and May, the SAMS decay phase
begins, as regions of heavy precipitation over the south-
ern Amazon and central Brazil decrease and gradually
migrate northwestward toward the equator and as the
rainy season along the eastern coast of northeast (NE)
Brazil gets under way and continues from April
through June (Rao and Hada 1990). Throughout the
decay phase of the SAMS, deep convection associated
with the Atlantic ITCZ is relatively weak.

3. Climate variability

a. Diurnal and mesoscale variability

1) DIURNAL CYCLE

The timing of daily maximum convection in the
Americas is location specific and closely related to to-
pographic features such as mountain ranges and coast
lines. In the NAMS there are large-scale shifts in the
regions of deep convection during the day from over
land during the afternoon and evening (especially along
the western slopes of the SMO) to offshore locations
during the morning hours (Figs. 5a,b; e.g., Higgins et al.
2003). In general, onshore flow begins during the morn-
ing transporting moist air from the Gulf of California
inland and up the slopes of the SMO. Enhanced by
inland mountain valley circulations, convection initiates
near midday along the western slopes and high ridges of
the SMO (Gochis et al. 2003; Fig. 6). During the eve-
ning and nighttime upslope flow reverses and generates
a westward-propagating zone of convergence that mi-
grates downslope across the coastal plains and out over

FIG. 4. Climatological mean accumulated precipitation (from
Xie and Arkin 1997) and vertically averaged climatological mean
moisture fluxes (from NCEP–NCAR reanalysis) for (a) DJF and
(b) June–August (JJA). Contour interval is 150 mm. Vector units
are kg (m s)�1. Orography higher than 1000 m is contoured in
black.

15 OCTOBER 2006 V E R A E T A L . 4981



FIG. 5. Evening (2100–2400 UTC) temporal frequency of cold clouds (infrared brightness temperature Tb � 235 K) during the (a)
boreal summer (JJA) and (c) austral summer (DJF). Pixel resolution is 0.5° � 0.5°. The period of analysis extends from 1983 to 1991.
(b), (d) Same as in (a) and (c), respectively, but for late night and early morning (0900–1200 UTC).

4982 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19

Fig 5 live 4/C



the Gulf of California by early morning. Convection, in
the form of small cloud systems, develops in coastal
zones of the narrow Isthmus of Central America during
the afternoon, and offshore nocturnal rainfall tends to
propagate away through the nighttime and morning
hours (Mapes et al. 2003b; Garreaud and Wallace
1997). The mechanism of propagation of the nocturnal
rainfall appears to involve gravity waves (Mapes et al.
2003a).

Because of interaction with the slope of the Central
American mountains, the diurnal cycle in Central
America is strongly dependent upon the magnitude of
the trade wind regime, which in turn is seasonally de-
pendent. The Pacific slope of Central America shows a
rainfall maximum occurring mainly in the afternoon
and early evening, while on the Caribbean side, the
diurnal cycle has a late evening–early morning precipi-
tation maxima that changes with altitude, with the
maximum occurring at altitudes between 600 and 2000
m in the early evening (Fernandez et al. 1996).

Satellite products have been used to document a co-
herent diurnal cycle of convective cloudiness over the
South American continent during the summer rainy
season [December–February (DJF)], and the peak is
observed at around 1800 LST (Garreaud and Wallace
1997; Negri et al. 1994; Sorooshian et al. 2002). During
the afternoon and evening convective clouds are par-
ticularly common over the Altiplano, along the north-
east coast of the continent, and in two parallel bands
over central Amazonia (Fig. 5c). The first two bands
arise from the concurrent timing of the thermodynamic
destabilization and the maximum strength of the plain-
to-mountain wind convergence over the central Andes
and land–sea breeze convergence along the coast, re-
spectively, while the latter have been interpreted as the
afternoon reactivation of inland-moving coastal squall
lines (e.g., Cohen et al. 1995; Garreaud and Wallace

1997). Marengo et al. (2005) found that the westerly
surface wind regime (related to enhanced SACZ) is
associated with a precipitation maximum in southwest
Amazonia during the afternoon (�1400 LST) followed
by a relatively weak maximum in the evening. During
the easterly regime, there is an early morning maximum
(�0200 LST) followed by a stronger maximum in the
afternoon (�1400 LST). Nighttime and early morning
convection tends to be more prevalent along the east-
ern slope of the central Andes and subtropical plains
(Fig. 5d; Nesbitt and Zipser 2003), and it can be as-
cribed to a decrease of the intensity of the compensat-
ing subsidence, which provides a fast linkage between
convectively active and inactive regions (Silva Dias et
al. 1987).

2) MESOSCALE VARIABILITY

While stability of the long-term precipitation clima-
tology over the core monsoon regions is provided by
stationary convective forcing mechanisms, regions pe-
ripheral to the core regions rely on mesoscale organi-
zation and propagation of convective storms or tran-
sient incursions of low-level moisture.

Transient disturbances including active or decaying
tropical storms from either the eastern Pacific or the
Gulf of Mexico, the passage of midlatitude troughs,
westward-propagating cyclonic perturbations in the
Tropics or various combinations of each of these phe-
nomena perturb specific components of the time mean
circulation thereby giving rise to increased mesoscale
variability in Central and North America. This variabil-
ity is manifested in phenomena such as pulsing of the
Gulf of California low-level jet (Douglas 1995; Berbery
and Fox-Rabinovitz 2003), intensification or reduction
of the land–sea and valley–mountain circulations (Ber-
bery 2001; Fawcett et al. 2002; Anderson et al. 2001)
and Gulf of California moisture surges (Douglas 1995;
Stensrud et al. 1997; Fuller and Stensrud 2000; Douglas
and Leal 2003; Higgins et al. 2004). Some climatological
features of long-lived, propagating, organized convec-
tion in the NAMS have been studied (e.g., McCollum et
al. 1995), but the specific mechanisms that promote me-
soscale organization and propagation remain elusive.

Velasco and Fritsch (1987) surveyed the mesoscale
convective complexes (MCCs) over South and North
America using Geostationary Operational Environ-
mental Satellite (GOES) infrared imagery and found a
large population of them over eastern South America
between 20° and 40°S, especially during December and
January. South American MCCs tend to develop dur-
ing the evening and reach their maximum size after
midnight, explaining the nocturnal maximum of con-
vective cloud frequency in this region (Fig. 5c). In com-

FIG. 6. Wet-day hourly rain rates for various elevation bands in
the SMO from the NAME Event Raingauge Network (NERN;
from Gochis et al. 2003).

15 OCTOBER 2006 V E R A E T A L . 4983



parison to North America, midlatitude South American
MCCs are significantly larger (average size of 500 000
km2 at the time of maximum extent) and last longer (12
h on average). A second cluster of MCCs is found over
the Gulf of Panama and along both slopes of the Co-
lombian Andes (see also Mapes et al. 2003b). In con-
trast, Carvalho and Jones (2001) found that most con-
vective systems (CSs) over Amazonia are relatively
small (average size of less than 100 000 km2), and short-
lived (3–6 h), and, therefore, do not meet MCC identi-
fication criteria. Tropical Rainfall Measuring Mission
(TRMM) satellite products reveal that wet season con-
vection over the Amazon region exhibits a vertical
structure intermediate to that observed over tropical
oceans (i.e., less vertically developed, less lighting) and
other continents (e.g., Nesbitt and Zipser 2003), al-
though intraseasonal variability is large (Petersen et al.
2002). Deep convection is also organized in small CSs
over the Altiplano (Garreaud 1999; Falvey and Gar-
reaud 2004).

One of the primary elements associated with the oc-
currence of subtropical South American CSs is the
SALLJ. National Centers for Environmental Predic-
tion (NCEP) and European Centre for Medium-Range
Weather Forecasts (ECMWF) reanalysis data suggest
that the strongest winds are between 15° and 20°S,
roughly over the Bolivian lowlands and Paraguay (Salio
et al. 2002; Marengo et al. 2004). Recently, pilot balloon
observations (Douglas et al. 1999) at Santa Cruz (18°S,
63°W, 30 km east of the Andean foothills) indicated
that, when present, the SALLJ exhibits maximum wind
speeds in excess of 15 m s�1 at about 1.5 km AGL and
a vertical shear above that maximum of 5 m s�1 km�1

(Douglas et al. 1999). The associated maximum of me-
ridional moisture flux occurs just to the east of the
Andes, between the surface and 700 hPa, sharply de-
caying eastward within 500 km of the foothills (Saulo et
al. 2000; Berbery and Collini 2000). Recently the South
American Low-Level Jet Experiment (SALLJEX) was
carried out in the region of sparse upper-air observa-
tions over central and subtropical South America dur-
ing the austral summer 2002/03 (Vera 2004). Prelimi-
nary results using SALLJEX observations describe the
significant diurnal variability of the SALLJ, with a
deep, late afternoon–evening maximum that might par-
tially explain the prevalence of nocturnal convection
south of 20°S (Nicolini et al. 2004; Zipser et al. 2004)

b. Synoptic variability

Tropical cyclones (TCs), which account for between
20% and 60% of the rainfall along the Pacific coast of
Mexico, also contribute as much as 25%–30% of the
seasonal rainfall at western interior locations (Engle-

hart and Douglas 2001), and they explain up to 15% of
summer precipitation over the coastal areas in the
United States (Larson et al. 2005). Tropical easterly
waves propagating into the NAM region from the tropi-
cal Atlantic can develop into a TC in the eastern Pacific
(Farfán and Zehnder 1997; Molinari et al. 2000). Ide-
alized modeling studies have shown that interactions
between the SMO, and easterly waves approaching
from the Gulf of Mexico, and the ITCZ can lead to
eastern Pacific cyclogenesis (Zehnder 1991; Zehnder et
al. 1999; Mozer and Zehnder 1996). Wave amplification
tends to occur in regions where the meridional gradient
in low-level potential vorticity reverses (Molinari et al.
1997; Molinari and Vollaro 2000). The ITCZ west of
the Central American coast can break down into mul-
tiple TCs because of a combined barotropic and baro-
clinic instability, dynamically supported by the reversal
in the meridional gradient of potential vorticity (Nieto-
Ferreira and Schubert 1997). Moisture surges, which
often traverse the length of the Gulf of California, ap-
pear to promote increased convective activity in north-
west (NW) Mexico and the southwest (SW) United
States and are related to the passage of tropical easterly
waves across western Mexico (Stensrud et al. 1997;
Fuller and Stensrud 2000).

In contrast to the numerous studies on easterly waves
in the tropical North Atlantic, there has not been as
much consideration of such disturbances in the tropical
South Atlantic and tropical South America. Easterly
waves from the Atlantic in combination with the tropi-
cal heat source in the western Amazon, and with me-
soscale features (sea-breeze and cloud-scale circula-
tions), promote squall lines that propagate over the
Amazon region (Cohen et al. 1995). Nevertheless,
while the daily precipitation distribution over the
northern part of South America during austral winter is
mainly due to the activity of easterly wave-type tropical
disturbances, daily precipitation variability over tropi-
cal South America during austral summer results from
the combined action of equatorial trades, westward-
traveling tropical disturbances, and the equatorward in-
cursions of midlatitude synoptic wave systems (Kayano
2003).

Synoptic-scale disturbances, which are particularly
large in the winter hemisphere but are also present dur-
ing summer, often reach sufficiently low latitudes to
affect the American monsoon systems (Fig. 7). Ander-
son and Roads (2002) investigated the characteristics of
summer precipitation variability in the southwestern
United States and found that midlatitude airmass intru-
sions interact with the monsoon ridge, thereby affecting
the location and timing of precipitation events near the
northern fringe of the NAMS. The synoptic patterns

4984 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19



associated with such shifts in the monsoon ridge, and
associated synoptic modulation of severe thunderstorm
activity, have been described by Maddox et al. (1995).

Over South America, the origin of the equatorward
extension of synoptic activity has been attributed to the
dynamical impact of the Andes cordillera on the syn-
optic-scale circulation. While upper-level synoptic sys-
tems tend to maintain their eastward motion as they
cross the Andes at low levels, they experience a marked
anticyclonic turning and equatorward elongation (e.g.,
Gan and Rao 1994; Vera et al. 2002). A schematic of
the evolution of surface low and high pressure systems
moving across South America shows that the persistent
subtropical anticyclone over the SE Pacific limits the
equatorward extension of low pressure systems ap-
proaching southwestern South America (Fig. 8a). As
the low pressure system moves inland, strong midlevel
westerly winds prevail over the subtropical Andes, re-
sulting in surface troughing east of the Andes (Fig. 8b).
The resulting northwestern Argentinean low (NAL; Se-
luchi et al. 2003) often merges with the monsoonal
Chaco Low (centered at about 20°S). The rather unique

thermal stratification of the NAL as well as its closed
circulation structure was exceptionally diagnosed dur-
ing SALLJEX, in agreement with previous modeling
studies (Saulo et al. 2004). The transient deepening of
the NAL favors the northerly transport of moist air to
the east of the Andes (Nieto Ferreira et al. 2003; Selu-
chi et al. 2003), feeding severe storms over the subtropi-
cal continental plains or mesoscale convective com-
plexes that contribute to further deepen the surface low
as it moves eastward into the Atlantic through the la-
tent heat released (Fig. 8c; Seluchi and Saulo 1998;
Vera et al. 2002). Following the cyclone, dry air surges
equatorward between the Andes and the Brazilian pla-
teau (Fig. 8d; e.g., Garreaud 1999) that can be accom-
panied by synoptic-scale bands of deep convection at
the leading edge, frequently associated with SACZ in-
tensification (e.g., Garreaud and Wallace 1998).

c. Intraseasonal variability

During the NH warm season, the Madden–Julian os-
cillation (MJO; e.g., Madden and Julian 1994) influ-

FIG. 7. Meridional heat flux by transient eddies at 925 hPa for (a) JJA and (b) DJF
(K m s�1). Black areas indicate terrain elevation in excess of 2000 m ASL.

15 OCTOBER 2006 V E R A E T A L . 4985



ences a number of different weather phenomena affect-
ing the North American monsoon. For example, there
is a strong relationship between the MJO and TC ac-
tivity in the Pacific and Atlantic Ocean basins (e.g.,
Liebmann et al. 1994; Maloney and Hartmann 2000).
The strongest tropical cyclones tend to develop when
the MJO favors enhanced precipitation and as the MJO
progresses eastward, the favored region for TC activity
also shifts eastward from the western Pacific to the east-
ern Pacific and finally to the Atlantic (Fig. 9; Higgins
and Shi 2001). While this relationship appears robust,
the MJO is one of many factors (e.g., SST conditions
and vertical wind shear) that contribute to the devel-
opment of TCs.

Evidence of active and quiescent periods of tropical
easterly wave and Gulf of California surge activity in
the NAM region motivated studies of relationships to
the MJO. Higgins et al. (2004) showed that the relative
location of the upper-level monsoon anticyclone in
midlatitudes at the time of the gulf surge affects the
response to the surge in the southwestern United States
(Figs. 10a,b). Wetter-than-normal (drier-than-normal)
conditions occur in the Southwest when the ridge axis
locates to the east (west) of the region. The response of
the ITCZ to the easterly wave forcing can result in gulf

surge events, but the response may depend on the
phase of the MJO. In particular, the westerly phase of
the MJO may result in larger vorticity in the vicinity of
the ITCZ, leading to a larger response (Zehnder et al.
1999).

The most distinctive pattern that characterizes in-
traseasonal rainfall variability over South America is a
dipolar one (Casarin and Kousky 1986; Nogués-Paegle
and Mo 1997). Enhanced precipitation over the SACZ
is accompanied by decreased rainfall in the subtropical
plains while the opposite phase is associated with in-
creased southward moisture flux from the Amazon re-
gion, and increased rainfall in the subtropical plains
(Figs. 10c,d; Diaz and Aceituno 2003). The dipole struc-
ture in precipitation is associated with distinct changes
in the position and intensity of the Bolivian High (Vera
and Vigliarolo 2000). Also, low-level zonal westerly
(easterly) winds over tropical Brazil during summer are
associated with an active (inactive) SACZ and net
moisture divergence (convergence) over SESA, imply-
ing a weak (strong) SALLJ (Herdies et al. 2002; Jones
and Carvalho 2002). Wave trains extending along the
South Pacific link convective pulses in the SPCZ and
SACZ regions (Grimm and Silva Dias 1995) and are
associated with both phases of the rainfall dipole pat-

FIG. 8. Schematic representation of the life cycle of surface low and high pressure systems moving across South America. The
symbols are as follows: Hs � subtropical high, Hm � migratory high, L � migratory low, and LT � lee trough.

4986 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19



tern (right panels of Figs. 10c,d; e.g., Liebmann et al.
2004a).

Carvalho et al. (2004) showed that the MJO plays a
role in modulating the persistence of intense SACZ
events while Liebmann et al. (2004a) found that rain
events in the SACZ (subtropical) region tend to occur
26 days prior to (2 days after) the peak in MJO con-
vection at 10°S, 110°W. Nevertheless, the dipole struc-
ture of convection anomalies not only exhibits variabil-
ity associated with the MJO, but also on time scales
around 22–28 days (Liebmann et al. 1999; Nogués-

Paegle et al. 2000), which implies that interaction
among different intraseasonal fluctuations might con-
trol variability in the dipole pattern.

Relative influences of ENSO and the MJO on the
monsoon must be separated and understood. Over the
NAM region, ENSO-related impacts are linked to rela-
tively broad north–south adjustments of the precipita-
tion patterns in the eastern Pacific, while MJO-related
impacts are linked to more regional north–south adjust-
ments and potential predictability increases in the core
monsoon region when both ENSO and the MJO are

FIG. 9. Composite evolution of 200-hPa velocity potential anomalies (106 m2 s�1) associated
with MJO events and points of origin of tropical disturbances that developed into hurricanes
or typhoons. Contour intervals are 1 � 106 m2 s�1, solid contours indicate positive values
(convection is suppressed), and dashed contours indicate negative values (convection is en-
hanced).

15 OCTOBER 2006 V E R A E T A L . 4987



taken into account (Higgins and Shi 2001). When the
seasonal ENSO phase is known, the explained variance
in summer precipitation is concentrated in a zonal band
near the ITCZ. However, if the seasonal MJO activity
is also known, then there is a plume of higher explained
variance that extends northward along the west coast of
Mexico and into the southwestern United States. Re-
cently Cazes-Boezio et al. (2003) found that ENSO
teleconnection patterns observed over the South Pacific
to South America in austral spring are due to changes
in the frequency of occurrence and amplitude of in-
traseasonal circulation regimes.

There is some evidence that MJO activity might in-

crease predictability over equatorial South America
(Wheeler and Weickmann 2001; Mo 2001). Intrasea-
sonal variations in the 10–90-day band may impact the
skill of numerical weather forecasts over South
America (Nogués-Paegle et al.1998) and the forecast
skill seems to increase (decrease) during periods of
strong (weak) convective activity associated with the
MJO (Jones and Schemm 2000).

d. Interannual variability

Interannual variability of the American monsoon
systems exhibits strong regional characteristics and
within each monsoon domain there are coherent re-

FIG. 10. Schematic of the typical 700-hPa circulation features of the NAMS (heights and winds) that accompany
(a) wet and (b) dry surges keyed to Yuma, AZ. Regional circulation anomalies characterizing periods of (c)
enhanced and (d) suppressed convective cloudiness over SESA during austral spring and summer. Letters H and
L and the associated circulation arrows represent anticyclonic and cyclonic circulation anomalies, respectively.
Here �T and �T stand for positive and negative temperature anomalies. The relative strength of the subtropical
jet stream is represented by arrows of different sizes. The arrow eastward from the central Andes indicates the
direction of the low-level wind anomaly. (From Diaz and Aceituno 2003.)

4988 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19



gional patterns in summer season rainfall anomalies
(Comrie and Glenn 1998; Gutzler 2004). Out-of-phase
relationships exist between precipitation in southwest-
ern North America and that in the Great Plains of the
United States (e.g., Higgins et al. 1999) and between
precipitation over the east-central Amazon and north-
east Brazil, and that over south Brazil, northeast Ar-
gentina, and Uruguay (Ropelewski and Halpert 1986;
Marengo 2002; Nogués-Paegle et al. 2002; Gan et al.
2004; Grimm 2003).

Previous studies have considered five factors for the
interannual variability of American monsoon precipi-
tation: SST anomalies, land surface conditions, the po-
sition and strength of the tropical convergence zones,
water vapor transport, and large-scale circulations. The
last two factors are largely controlled by internal atmo-
spheric dynamics. However seasonal anomalies in the
large-scale circulation may be predictable to the extent
that they are related to the boundary conditions. The
fraction of total interannual variance explained by any
one of the factors above is small, suggesting that pre-
dictability may be modest or that all of these factors
(which actually are not independent) must be consid-
ered together, though their relative influences on pre-
dictability must be understood (Koster et al. 2000;
Marengo et al. 2003).

1) ROLE OF SST ANOMALIES

SST effects on the monsoons can be both local [e.g.,
Gulf of California (GC), Gulf of Mexico (GM), Peru-
vian coast, and the southwestern Atlantic] and/or re-
mote (e.g., tropical Pacific or tropical Atlantic), and in
many cases the influences are modest. Carleton et al.
(1990) showed that monsoon precipitation in the south-
west United States is negatively correlated with SST
anomalies along the northern Baja coast, while Huang
and Lai (1998) found positive correlations with SST
anomalies over the Gulf of Mexico. Roughly 80% of
the rainfall in the Arizona/New Mexico region occurs
after SST values in the northern Gulf of California ex-
ceed 28.5°C, which seems to be important for the tim-
ing, intensity, and extent of the monsoon (Mitchell et al.
2002). SST anomalies near the Peruvian coast may af-
fect rainfall over South America by changing the land–
sea temperature contrast (Yu and Mechoso 1999).

Remote SST anomalies in the equatorial Pacific as-
sociated with El Niño (La Niña) tend to be associated
with anomalously dry (wet) North American summer
monsoons in the core monsoon region over western and
southern Mexico (e.g., Gutzler and Preston 1997; Hig-
gins et al.1999; Hu and Feng 2002). Rainfall over the
east-central Amazon and northeast Brazil (southeast-
ern South America and central Chile) tends to be below

(above) normal during the warm phase of ENSO (e.g.,
Moura and Shukla 1981; Pisciottano et al. 1994; Grimm
et al.1998).

Tropical Atlantic SST anomalies have a strong influ-
ence on precipitation in tropical South America (e.g.,
Mechoso et al. 1990; Giannini et al. 2001, and refer-
ences therein) and in particular the gradient of SST
anomalies between the tropical North and South At-
lantic appears to be the key element associated with
rainfall anomalies over northeast Brazil and north-
central Amazonia (Marengo 1992; Hastenrath and
Greischar 1993; Uvo et al. 1998). Summer precipitation
anomalies are correlated with SST anomalies in the
western subtropical South Atlantic (WSSA; e.g., Diaz
et al. 1998). Positive (negative) SST anomalies in the
WSSA region are associated with moisture transport
from tropical latitudes southeastward (eastward), and
with positive rainfall anomalies over northeastern Ar-
gentina and southern Brazil (SACZ region; Doyle and
Barros 2002; Robertson and Mechoso 2003).

2) ROLE OF LAND SURFACE CONDITIONS

Gutzler and Preston (1997) and Gutzler (2000) found
that heavy snow in the southern Rockies tends to be
associated with deficient summer rainfall in portions of
the southwest United States and vice versa in analogy
with the results of Yang et al. (1996) for the Asian
monsoon. Higgins et al. (1998) argued that the effects
of SST anomalies on cold season precipitation indi-
rectly affects warm season rainfall, since they play a
role in determining the initial springtime soil moisture
conditions and vegetative cover, which in turn can feed
back upon the climate during the warm months.

Modeling studies have demonstrated that regions
such as southern Amazonia and the SAMS core regions
exhibit relatively low climate predictability on seasonal-
to-interannual time scales, since circulation and rainfall
anomalies in these regions are more dependent on re-
gional forcing than remote forcing (Goddard et al.
2003; Koster et al. 2000; Marengo et al. 2003). The
SAMS onset is characterized by the transition of a pol-
luted atmosphere (due to winter subsidence conditions
and biomass burning) to a clear one during the rainy
season. Biogenic aerosol and aerosol produced by bio-
mass burning have a direct influence on the surface and
tropospheric energy budget because of their capacity to
scatter and absorb solar radiation (Artaxo et al. 1990).
Recent modeling and observational results have indi-
cated that the aerosol plume produced by biomass
burning at the end of the dry season is transported to
the south, where it may interact with frontal systems
and possibly impact the precipitation regime (Freitas et

15 OCTOBER 2006 V E R A E T A L . 4989



al. 2005) through the radiative forcing of cloud micro-
physical processes (Silva Dias et al. 2002).

3) ROLE OF TROPICAL CONVERGENCE ZONES

ENSO-related impacts on precipitation in the NAMS
are strongly linked to meridional adjustments of the
ITCZ (e.g., Yu and Wallace 2000; Higgins and Shi
2001). During El Niño the ITCZ shifts southward, re-
sulting in an anomalously strong local meridional (Had-
ley) circulation and a reduction in seasonal mean rain-
fall over much of Mexico and the Caribbean (Higgins
and Shi 2001; Hu and Feng 2002). During La Niña the
ITCZ shifts northward, with attendant weakening of
the local Hadley circulation and an increase in NAMS
precipitation.

During austral summer, SST anomalies in the tropi-
cal Atlantic modulate the location of the ITCZ and thus
precipitation over northeast Brazil (e.g., Nogués-Paegle
and Mo 2002). Over the southwestern Atlantic, a dipo-
lar structure of the SST anomalies accompanies inter-
annual variations of the SACZ (Robertson and
Mechoso 2000). Coupled atmosphere–ocean GCM
simulations indicate that warm SST anomalies in the
South Atlantic lead to a northward displacement and
intensification of the SACZ. An intensified SACZ
tends to induce underlying cooler SST anomalies
(Chavez and Nobre 2004). Enhanced convective activ-
ity in the SACZ may lead to intensification of the Eur-
asian teleconnection pattern (Grimm and Silva Dias
1995).

4) ROLE OF MOISTURE TRANSPORT

The Gulf of California and Great Plains low-level jets
contribute to the summer precipitation and hydrology
in the southwestern United States and Mexico and cen-
tral United States, respectively (e.g., Bosilovich and
Schubert 2002). In South America, the SALLJ east of
the Andes shows a weak tendency to be stronger and
more frequent when SST anomalies are positive in the
tropical Pacific (Marengo et al. 2004). Also, numerical
simulations and diagnostic studies have shown that the
El Niño of 1998 featured more frequent and intense
summer SALLJ episodes than the La Niña of 1999
(Douglas et al. 1999; Saulo et al. 2000). Nevertheless,
there is considerable interannual variability of these
LLJs and water vapor transport that is quite indepen-
dent of ENSO (e.g., Hu and Feng 2001; Berbery and
Barros 2002).

5) ROLE OF THE LARGE-SCALE CIRCULATION

Interannual variations in large-scale teleconnection
patterns, such as the Pacific–North American (PNA)

pattern, are related to interannual variations in key
components of the North American monsoon (Carle-
ton et al. 1990), including the upper-tropospheric mon-
soon anticyclone over southwestern North America
(Higgins et al. 1998). While some of this variability is
directly linked to ENSO, at least a portion of it is re-
lated to the leading patterns of extratropical climate
variability (e.g., the Arctic Oscillation), which are
largely independent of ENSO (e.g., Higgins et al. 2000).

During warm (cold) ENSO phases, Rossby wave
trains emanating from either the western or central Pa-
cific toward SH high latitudes [Pacific–South America
(PSA) patterns] maintain a distinctive cyclonic (anticy-
clonic) circulation anomaly in the vicinity of South
America that enhances (suppresses) precipitation over
SESA (e.g., Grimm et al. 2000). Despite the fact that
the ENSO signal projects strongly on the PSA patterns,
numerical simulations show that they can also be gen-
erated by internal atmospheric dynamics (Cai and
Watterson 2002). Recently Vera et al. (2004) showed
that differences in the teleconnection patterns in the
SH among different ENSO warm events are related to
changes in the intensity and location of the SPCZ as
well as to differences in the activity of the three leading
modes of circulation variability in the SH, independent
of ENSO. Silvestri and Vera (2003) have shown that the
Antarctic Oscillation, which is largely independent of
ENSO, influences interannual precipitation changes
over subtropical South America (especially during aus-
tral winter and spring).

e. Variability on decadal (and longer) time scales

The summer monsoon in the southwest United States
seems to be modulated by longer-term (decade scale)
fluctuations in the North Pacific SSTs associated with
the Pacific decadal oscillation (PDO; Higgins and Shi
2000). The link between the North Pacific wintertime
SST pattern and the summer monsoon appears to be
established via the impact of variations in the Pacific jet
on west coast precipitation regimes during the preced-
ing winter. This affects local land-based sources of
memory in the southwestern United States, which in
turn influence the subsequent timing and intensity of
the summer monsoon (e.g., Castro et al. 2001).

A number of studies have reported the existence of
decadal and longer time-scale variability in South
American rainfall, related to ocean surface changes on
those time scales in both the Pacific and the Atlantic
Oceans (Robertson and Mechoso 2000; Zhou and Lau
2001). Interdecadal variability of around a 15-yr period
has been observed in SACZ activity, in southwest At-
lantic SSTs and river flows of the La Plata Basin (Rob-
ertson and Mechoso 2000). Positive correlations have

4990 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19



been found between SST trends in the tropical South
Atlantic (Fig. 11c) and precipitation anomalies in
northeast Brazil (Wagner 1996), Amazonia (Marengo
2004)) and SESA (Fig. 11a; Liebmann et al. 2004b).
Marengo et al. (2004) showed positive trends in SALLJ
activity (Fig. 12a) consistent with trends in the North
Atlantic Oscillation (NAO; Fig. 12d) and the PDO
(Fig. 12e). In particular, the climate shift that occurred
in the Pacific around the late 1970s (Mantua et al. 1997)
is significantly related with an increase of SALLJ ac-
tivity (Marengo 2004) as well as with an increase in
precipitation over SESA.

4. Land surface variations in the American
monsoon regions

Land use/land cover changes are occurring in the
core region of the both NAMS and SAMS regions.
Overgrazing within the steep, semiarid terrain of the
SMO has resulted in dramatic changes in surface veg-
etation and soil characteristics, which, in turn, produce
several important hydrological responses that have not
yet been assessed (Viramontes and Descroix 2003).
There also exists, along the western slope of the SMO,
large expanses of deciduous tropical forests, which ex-
perience a dramatic green-up cycle immediately follow-
ing the onset of the monsoon. It has been suggested that
following this green-up these large expanses of forest
may serve as critical pathways for the recycling of mon-
soon precipitation back to the atmosphere, although
the effects of such activities on the regional climate are
still open to question.

The Amazon basin is the world’s most extensive
tropical rain forest and more than one-third of the
world’s woody biomass is located in South America,
with 27% in Brazil alone. However, the rate of defor-
estation in Central and South America is among the
highest in the world at an annual average rate of 0.48%.
From the last 30 yr to present, the conversion of forest
land to other uses (such as agricultural activities) has
been causing the deforestation and degradation of the
forest ecosystem. Since the mid-1980s, atmospheric

→

FIG. 11. JFM rainfall (mm season�1) over southern Brazil from
station data vs year. Filled circles represent years 1948–75. Filled
squares represent years 1976–99. Dashed curves represent linear,
least squares fits for segments 1948–75, 1976–99, and 1948–99. (b)
Same as in (a) except ordinate is average JFM discharge at Salto
Caxias, Brazil, on the Iguazú River, near the confluence with the
Paraná. (c) Same as in (a) except that ordinate is an index of
average SST (21.9°–27.6°S on a Gaussian grid, 28.2°–41.2°W; from
Liebmann et al. 2004b).

15 OCTOBER 2006 V E R A E T A L . 4991



GCM simulations have suggested that deforestation in
tropical South America can cause a significant decrease
in precipitation and evapotranspiration and a decrease
in moisture convergence, especially in central and
northern Amazonia (e.g., Marengo and Nobre 2001).
New developments in dynamic vegetation schemes and
coupled climate–carbon models (Cox et al. 2000; Betts
et al. 2004; Huntingford et al. 2004) have shown that the
physiological forcing of stomatal closure can contribute
20% to the rainfall reduction in the Amazon associated
with rising atmospheric CO2 levels. Nevertheless, the
influence of land surface variations in explaining the
slightly positive rainfall trends documented in southern
Amazonia since the mid-1970s (Marengo 2004) and the
significant positive trends detected in rainfall and
streamflow over the subtropical portion of the La Plata
Basin (Figs. 11b,c; Genta et al. 1998; Liebmann et al.
2004b) is still unclear. It has been suggested that an
increase in southern Amazonia rainfall could be due to
mesoscale circulations triggered by the deforestation
(Silva Dias et al. 2002).

5. Monsoon hydrology

From the standpoint of land surface hydrology, mon-
soon regions and systems have several characteristics

relevant for the land surface hydrology. A comparison
between the mean monthly runoff as a fraction of the
annual total precipitation (Fig. 13) and the mean
monthly runoff as a fraction of average annual total
precipitation (Fig. 14) shows that runoff is a more im-
portant process in wet regions. The average annual run-
off ratio for the Little Colorado River (NAMS) is only
about 0.1, while that for the Mahaweli Ganga (India
region) is around 0.5. Furthermore, the relatively mod-
est winter (nonmonsoon) precipitation in the NAMS
leads to considerably higher runoff, on average, than
does the monsoon precipitation, primarily because of
reduced evapotranspiration in winter, and the contri-
bution within small portions of the NAMS domain
(e.g., the Little Colorado River) of winter snow that
contributes disproportionately to winter and spring
runoff. The figures also show that there are strong dif-
ferences in the seasonality of runoff and streamflow
across the regions as well. In the Sahel, precipitation
and runoff are very low in months outside the monsoon
period, whereas in India, there is significant runoff out-
side the monsoon period.

In the SAMS the seasonal signature of runoff is ex-
tremely damped and is only modestly higher in the
months following the precipitation peak. The rivers in
the southern Amazon basin, extended over the north-
ern part of the SAMS core region, exhibit their maxi-
mum discharges around 2–3 months after the rainfall
peaks, while those in the northern Amazon basin are
more sensitive to the interannual variability linked to
strong ENSO events (Marengo 2004). On the other
hand, the upper Paraná basin located over the central
and southern portion of the SAMS region exhibits
maximum discharges during February and March asso-
ciated with copious summer rainfall in the SACZ
(Camilloni and Barros 2000).

Therefore, there are important similarities in the
land surface hydrology of most of the monsoon regions
(strongly seasonal runoff, but with peak runoff ratios
much lower than in a strongly seasonal nonmonsoon
environment, primarily due to high evaporative de-
mand during the maximum precipitation season). How-
ever, there are major differences as well. In the arid to
semiarid monsoon environments (NAMS and Sahel)
runoff ratios are much lower than in the humid SAMS
and India regions, which also have more uniform runoff
over the year. Again, while these differences are attrib-
utable in considerable part to the characteristics of the
land surface forcing (especially precipitation), some of
the differences, especially in the seasonality and tem-
poral variability of runoff, are due to catchment-
specific topographic, geological, and vegetation charac-
teristics.

FIG. 12. Normalized annual departures of (a) SALLJ event an-
nual counts, (b) Southern Oscillation index (SOI), (c) El Niño-3.4
SSTs, (d) NAO, and (e) PDO indexes. Thick lines represent the
11-yr moving averages.

4992 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19



6. Future challenges

Despite significant advances in our understanding of
the American monsoon systems, many challenges re-
main. At local scales, additional attention to the low-
level circulation and precipitation patterns is needed.
Some principal research questions in common to both
the NAMS and the SAMS include the following: 1)
What are the relationships between local low-level cir-
culation features (e.g., the low-level jets; mountain–
valley circulation) and the diurnal cycle of moisture and
convection? 2) What are the dominant sources of pre-
cipitable moisture for monsoon precipitation? 3) What
are the relative roles of local variations in sea surface
temperature and land surface parameters (topography,
soil moisture, and vegetation cover) in modulating pre-
cipitation?

The diurnal cycle (and related processes and feed-
backs) is poorly represented in models not only over
the core monsoon regions, but also at most locations
over the Americas during the warm season. The various
atmospheric pathways for moisture transport into the
Americas should be identified and understood. An im-
proved characterization of deep convection will play a
critical role in accurate depiction of moisture transport
over the continents. The relative contributions of local
oceanic forcing, soil moisture, and vegetation to warm

season precipitation variability in the monsoon regions
are not known. At present, we know relatively little
about land surface conditions and fluxes over most of
the NAMS and SAMS regions, and characterizing the
surface fluxes is critical for understanding and simulat-
ing the hydrologic cycle. The role of the transport of
aerosols produced by biomass burning in the Amazon
region on the SAMS variability should also be further
explored and basic questions that need further investi-
gation are as follows: What is the comparative impact of
remote climate forcing and of the regional effects of
biomass burning generated aerosol in the beginning of
the rainy season? Is the local effect of aerosol more
significant through a radiative impact or through cloud
microphysics?

On regional scales, an improved description and un-
derstanding of intraseasonal aspects of the monsoon,
especially those related to the MJO is needed. Some
general questions include the following: 1) What is the
nature of the relationship between the MJO and key
components of the monsoon systems (e.g., low-level
jets, gulf surges, etc.)? 2) What are the relationships
between the MJO and extreme weather events in the
Americas? 3) What are the relative influences of the
MJO and ENSO on monsoon precipitation?

Of practical significance is the fact that there is a

FIG. 13. Normalized (by annual average) mean monthly runoff for selected rivers in each
basin. NAMS [river (Little Colorado near Cameron, AZ), drainage area (68 634 km2), (lat,
lon) � (35.92°N, 111.57°W)]; SAMS [river (Rio Grande de Bahia near Boqueirao, Brazil),
drainage area (65 900 km2), (lat, lon) � (11.33°S, 43.80°W)]; India [river (Mahaweli Ganga
near Manampitiya, Sri Lanka), drainage area (7343 km2), (lat, lon) � (7.92°N, 81.08°E)]; Sahel
[river (Senegal River near Bakel, Senegal), drainage area (218 000 km2), (lat, lon) � (14.90°N,
12.45°W)]; PNW [river (Chehalis River at Porter, WA), drainage area (3351 km2), (lat, lon)
� (46.94°N, 123.31°W)].

15 OCTOBER 2006 V E R A E T A L . 4993

Fig 13 live 4/C



coherent relationship between the phase of the MJO
and the precipitation pattern around the global Tropics.
Additional studies are needed to determine how the
phase of the MJO relates to the frequency and intensity
of hurricanes and tropical storms, what fraction of the
summer rainfall over the Americas is due to tropical
storms and hurricanes, and whether interannual-to-
interdecadal variability of the MJO contributes to
tropical cyclone frequency and intensity. While ENSO-
related impacts on the monsoons are reasonably well
documented, MJO-related impacts, as well as the rela-
tive influences of the MJO and ENSO, are not well
understood.

At continental scales, the focus should be on im-
proved description and understanding of the spatial and
temporal linkages between warm season precipitation,
the large-scale circulation patterns, and the dominant
boundary forcing parameters (both land and ocean).
Some general scientific questions are as follows: 1)
How is the evolution of monsoonal precipitation re-
lated to the seasonal evolution of the oceanic and con-
tinental boundary conditions? 2) What are the relation-
ships between interannual variations in the boundary
conditions, the atmospheric circulation, and the conti-
nental hydrologic regime? 3) What is the correlation
between the anomaly sustaining atmospheric circula-
tion and the land and ocean surface boundary condi-
tions that characterize precipitation and temperature
anomalies during the summer?

Studies that enhance our dynamical understanding of

the seasonal march of the monsoon and its variability
are critical. The relative importance of the land and
ocean influences on monsoon precipitation changes
with the seasons as does the potential predictability of
circulation and precipitation anomalies. The remote in-
fluences of other low-frequency modes of variability
besides ENSO (e.g., PDO) on the warm season precipi-
tation regimes over the Americas may be important.
There is evidence that potentially predictable anoma-
lies of soil moisture, snow cover, and vegetation play a
role in the seasonal variability of precipitation patterns,
and it has been suggested that there are important feed-
backs (of either sign) between the atmosphere and land
surface. Prediction of the detailed distribution of con-
tinental precipitation is a challenging task since it re-
quires the skillful modeling of the subtle interplay be-
tween land surface and oceanic influences such as the
complicating influences of terrain and coastal geom-
etry.

The availability of comprehensive datasets that docu-
ment key features of the American monsoons is crucial.
Depiction of the diurnal cycle of precipitation and at-
mospheric circulation over the NAMS and SAMS re-
gions and long-term monitoring of oceanic and land
surface conditions are but two examples of the types of
datasets that are needed.

MESA and NAME

The WCRP/CLIVAR/VAMOS Program is imple-
menting two international monsoon programs in the

FIG. 14. Monthly runoff ratios (monthly average streamflow as a fraction of average of
annual precipitation) for five river basins listed in the caption of Fig. 13.

4994 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19

Fig 14 live 4/C



Americas: the Monsoon Experiment South America
(MESA) and the North American Monsoon Experi-
ment (NAME), aimed at improving warm season
precipitation forecasts over the Americas on seasonal-
to-interannual time scales. To help MESA and NAME
achieve this goal, CLIVAR and the Global Energy
and Water Cycle Experiment (GEWEX) are imple-
menting empirical and modeling studies plus dataset
development and enhanced monitoring activities, and
several field campaigns [e.g., SALLJEX, NAME 2004,
and La Plata Basin Program (PLATIN)]. As a conse-
quence, these programs will deliver improved infra-
structure to monitor and predict the American mon-
soons, more comprehensive understanding of climate
variability and predictability, strengthened multina-
tional scientific collaboration across Pan America, and
improved climate models that predict monsoon vari-
ability months to seasons in advance.

To meet these challenges, the VAMOS Program is
developing a modeling strategy that incorporates both
MESA and NAME activities. The strategy includes
baseline seasonal simulations that correspond to major
field campaigns and multiyear simulations focused on
key physical processes (e.g., the diurnal cycle of con-
vection). Preliminary SALLJEX modeling coordinated
activities, which included participation by groups from
South America and the United States, have provided
useful hints to understanding and improving the poor
model performances in simulating precipitation associ-
ated with the MCSs over central South America
(Paegle et al. 2004). The North American Monsoon
Assessment Project (NAMAP; Gutzler et al. 2004)
proved to be an excellent way to benchmark and assess
current global and regional model simulations of the
North American monsoon. In addition, joint experi-
mental prediction (e.g., sensitivity to soil moisture and
SST) and product development (e.g., hazards assess-
ments and drought monitors) activities are in the plan-
ning stages. Additional information about MESA and
NAME is available on the VAMOS Web site (http://www.
clivar.org/organization/vamos/vamos.php; VAMOS
datasets are available at http://www.joss.ucar.edu/
vamos/data/).

Acknowledgments. The coauthors take great plea-
sure in acknowledging the VAMOS community for the
tremendous scientific contributions to our understand-
ing of the American monsoon systems as summarized
here. A very special thanks is extended to Valery De-
temmermann (WCRP), Michael Patterson (NOAA/
OGP), Carlos Ereño, Andreas Villwock (ICPO), Gus
Emanuel, Steven Williams, and Jose Meitin (UCAR/
JOSS) for outstanding support. Finally, we thank the

reviewers, E. Hugo Berbery and Todd Mitchell, for
valuable comments that substantially improved the
manuscript.

REFERENCES

Adams, D. K., and A. C. Comrie, 1997: The North American
monsoon. Bull. Amer. Meteor. Soc., 78, 2197–2213.

Adang, T. C., and R. L. Gall, 1989: Structure and dynamics of the
Arizona monsoon boundary. Mon. Wea. Rev., 117, 1423–
1438.

Anderson, B. T., and J. O. Roads, 2002: Regional simulation of
summertime precipitation over the southwestern United
States. J. Climate, 15, 3321–3342.

——, ——, S.-C. Chen, and H.-M. H. Juang, 2001: Model dynam-
ics of summertime low-level jets over northwestern Mexico.
J. Geophys. Res., 106, 3401–3413.

Artaxo, P., W. Maenhaut, H. Storms, and R. Van Grieken, 1990:
Aerosol characteristics and sources for the Amazon Basin
during the wet season. J. Geophys. Res., 95, 16 971–16 985.

Berbery, E. H., 2001: Mesoscale moisture analysis of the North
American monsoon. J. Climate, 14, 121–137.

——, and E. Collini, 2000: Springtime precipitation and water
vapor flux convergence over southeastern South America.
Mon. Wea. Rev., 128, 1328–1346.

——, and V. R. Barros, 2002: The hydrologic cycle of the La Plata
basin in South America. J. Hydrometeor., 3, 630–645.

——, and M. S. Fox-Rabinovitz, 2003: Multiscale diagnosis of the
North American monsoon system using a variable-resolution
GCM. J. Climate, 16, 1929–1947.

Betts, R. A., P. M. Cox, M. Collins, P. Harris, C. Huntingford, and
C. D. Jones, 2004: The role of ecosystem-atmosphere inter-
actions in simulated Amazonian precipitation decrease and
forest dieback under global change warming. Theor. Appl.
Climatol., 78, 157–175.

Bosilovich, M. G., and S. D. Schubert, 2002: Water vapor tracers
as diagnostics of the regional hydrologic cycle. J. Hydro-
meteor., 3, 149–165.

Byerle, L. A., and J. Paegle, 2002: Description of the seasonal
cycle of low-level flows flanking the Andes and their inter-
annual bariability. Meteorologica, 27, 71–88.

Cai, W., and I. G. Watterson, 2002: Modes of interannual vari-
ability of the Southern Hemisphere circulation simulated by
the CSIRO climate model. J. Climate, 15, 1159–1174.

Camilloni, I., and V. R. Barros, 2000: The Paraná River response
to El Niño 1982–83 and 1997–98 events. J. Hydrometeor., 1,
412–430.

Campetella, C., and C. Vera, 2002: The influence of the Andes
Mountains on the South American low-level flow. Geophys.
Res. Lett., 29, 1826, doi:10.1029/2002GL015451.

Carleton, A. M., 1987: Summer circulation climate of the Ameri-
can Southwest, 1945–1984. Ann. Assoc. Amer. Geogr., 77,
619–634.

——, D. A. Carpenter, and P. J. Weser, 1990: Mechanisms of in-
terannual variability of southwest United States summer pre-
cipitation maximum. J. Climate, 3, 999–1015.

Carvalho, L. M. V., and C. Jones, 2001: A satellite method to
identify structural properties of mesoscale convective systems
based on the maximum spatial correlation tracking technique
(MASCOTTE). J. Appl. Meteor., 40, 1683–1701.

15 OCTOBER 2006 V E R A E T A L . 4995



——, ——, and B. Liebmann, 2004: The South Atlantic conver-
gence zone: Intensity, form, persistence, and relationships
with intraseasonal to interannual activity and extreme rain-
fall. J. Climate, 17, 88–108.

Casarin, D. P., and V. E. Kousky, 1986: Anomalias de precipi-
tação no sul do Brasil e variações na circulação atmosférica.
Rev. Bras. Meteor., 1, 83–90.

Castro, C. L., T. B. McKee, and R. A. Pielke Sr., 2001: The rela-
tionship of the North American monsoon to tropical and
North Pacific Sea surface temperatures as revealed by obser-
vational analyses. J. Climate, 14, 4449–4473.

Cazes-Boezio, G., A. Robertson, and C. R. Mechoso, 2003: Sea-
sonal dependence of ENSO teleconnections over South
America and relationships with precipitation in Uruguay. J.
Climate, 16, 1159–1176.

Chaves, R. R., and P. Nobre, 2004: Interactions between sea sur-
face temperature over the South Atlantic Ocean and the
South Atlantic Convergence Zone. Geophys. Res. Lett., 31,
L03204, doi:10.1029/2003GL018647.

Chou, C., and J. D. Neelin, 2003: Mechanisms limiting the north-
ward extent of the northern summer monsoons over North
America, Asia, and Africa. J. Climate, 16, 406–425.

Cohen, J., M. Silva Dias, and C. Nobre, 1995: Environmental
conditions associated with Amazonian squall lines: A case
study. Mon. Wea. Rev., 123, 3163–3174.

Comrie, A. C., and E. C. Glenn, 1998: Principal components-
based regionalization of precipitation regimes across the
southwest United States and northern Mexico, with an appli-
cation to monsoon precipitation variability. Climate Res., 10,
201–215.

Cox, P. M., R. A. Betts, C. D. Jones, S. A. Spall, and I. J. Totter-
dell, 2000: Acceleration of global warming due to carbon-
cycle feedbacks in a coupled climate model. Nature, 408, 184–
187.

Diaz, A. F., and P. Aceituno, 2003: Atmospheric circulation
anomalies during episodes of enhanced and reduced convec-
tive cloudiness over Uruguay. J. Climate, 16, 3171–3185.

——, C. D. Studzinski, and C. R. Mechoso, 1998: Relationships
between precipitation anomalies in Uruguay and southern
Brazil and sea surface temperature in the Pacific and Atlantic
Oceans. J. Climate, 11, 251–271.

Douglas, M. W., 1995: The summertime low-level jet over the
Gulf of California. Mon. Wea. Rev., 123, 2334–2347.

——, and J.-C. Leal, 2003: Summertime surges over the Gulf of
California: Aspects of their climatology, mean structure and
evolution from radiosonde, NCEP reanalysis, and rainfall
data. Wea. Forecasting, 18, 55–74.

——, R. A. Maddox, K. Howard, and S. Reyes, 1993: The Mexi-
can monsoon. J. Climate, 6, 1665–1667.

——, M. Nicolini, and C. Saulo, 1999: The low level jet at Santa
Cruz, Bolivia, during the January–March 1998, pilot balloon
observations and model comparison. Preprints, 10th Conf.
on Global Change Studies, Dallas, TX, Amer. Meteor. Soc.,
11–12.

Doyle, M. E., and V. R. Barros, 2002: Midsummer low-level cir-
culation and precipitation in subtropical South America and
related sea surface temperature anomalies in the South At-
lantic. J. Climate, 15, 3394–3410.

Englehart, P. J., and A. V. Douglas, 2001: The role of Eastern
North Pacific tropical storms in the rainfall climatology of
western Mexico. Int. J. Climatol., 21, 1357–1370.

Falvey, M., and R. Garreaud, 2004: Variability of moisture and
convection over the Central Andes during SALLJEX.
CLIVAR Exchanges, Vol. 9, No. 1, International CLIVAR
Project Office, Southampton, United Kingdom, 11–13.

Farfán, L. M., and J. A. Zehnder, 1997: Orographic influence on
the synoptic-scale circulations associated with the genesis of
Hurricane Guillermo. Mon. Wea. Rev., 125, 2683–2698.

Fawcett, P. J., J. R. Stalker, and D. S. Gutzler, 2002: Multistage
moisture transport into the interior of northern Mexico dur-
ing the North American summer monsoon. Geophys. Res.
Lett., 29, 2094, doi:10.1029/2002Gl015693.

Fernandez, W., R. E. Chacon, and J. W. Melgarejo, 1996: On the
rainfall distribution with altitude over Costa Rica. Geofisica,
44, 57–72.

Figueroa, S. N., and C. Nobre, 1990: Precipitation distribution
over central and western tropical South America. Climana-
lise, 5, 36–44.

Freitas, S. R., K. M. Longo, M. A. F. Silva Dias, P. L. Silva Dias,
F. S. Recuero, R. Chatfield, E. Prins, and P. Artaxo, 2005:
Monitoring the transport of biomass burning emissions in
South America. Environ. Fluid Mech., 5, 135–167.

Fu, R., B. Zhu, and R. E. Dickinson, 1999: How do atmosphere
and land surface influence seasonal changes of convection in
the tropical Amazon? J. Climate, 12, 1306–1321.

Fuller, R. D., and D. J. Stensrud, 2000: The relationship between
tropical easterly waves and surges over the Gulf of California
during the North American monsoon. Mon. Wea. Rev., 128,
2983–2989.

Gan, M. A., and V. B. Rao, 1994: The influence of the Andes
Cordillera on transient disturbances. Mon. Wea. Rev., 122,
1141–1157.

——, V. E. Kousky, and C. F. Ropelewski, 2004: The South
America monsoon circulation and its relationship to rainfall
over west-central Brazil. J. Climate, 17, 47–66.

Gandu, A. W., and P. L. Silva Dias, 1998: Impact of tropical heat
sources on the South American tropospheric upper circula-
tion and subsidence. J. Geophys. Res., 103, 6001–6015.

Garreaud, R. D., 1999: Multiscale analysis of the summertime pre-
cipitation over the central Andes. Mon. Wea. Rev., 127, 901–
921.

——, and J. M. Wallace, 1997: The diurnal march of convective
cloudiness over the Americas. Mon. Wea. Rev., 125, 3157–
3171.

——, and ——, 1998: Summertime incursions of midlatitude air
into subtropical and tropical South America. Mon. Wea. Rev.,
126, 2713–2733.

Genta, J. L., G. Perez-Iribarren, and C. R. Mechoso, 1998: A re-
cent increasing trend in the streamflow of rivers in southeast-
ern South America. J. Climate, 11, 2858–2862.

Giannini, A., J. C. H. Chiang, M. Cane, Y. Kushnir, and R. Seager,
2001: The ENSO teleconnection to the tropical Atlantic
Ocean: Contributions of the remote and local SSTs to rainfall
variability in the tropical Americas. J. Climate, 14, 4530–4544.

Gochis, D. J., J.-C. Leal, C. J. Watts, W. J. Shuttleworth, and
J. Garatuza-Payan, 2003: Preliminary diagnostics from a new
event-based precipitation observation network in support of
the North American Monsoon Experiment (NAME). J. Hy-
drometeor., 4, 974–981.

Goddard, L., A. G. Barnston, and S. J. Mason, 2003: Evaluation of
the IRI’S “net assessment” seasonal climate forecasts: 1997–
2001. Bull. Amer. Meteor. Soc., 84, 1761–1781.

4996 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19



Grimm, A. M., 2003: The El Niño impact on the summer monsoon
in Brazil: Regional processes versus remote influences. J. Cli-
mate, 16, 263–280.

——, and P. L. Silva Dias, 1995: Analysis of tropical–extratropical
interactions with influence functions of a barotropic model. J.
Atmos. Sci., 52, 3538–3555.

——, S. Ferraz, and J. Gomez, 1998: Precipitation anomalies in
southern Brazil associated with El Niño and La Niña events.
J. Climate, 11, 2863–2880.

——, V. R. Barros, and M. E. Doyle, 2000: Climate variability in
southern South America associated with El Niño and La
Niña events. J. Climate, 13, 35–58.

Gutzler, D. S., 2000: Covariability of spring snowpack and sum-
mer rainfall across the southwest United States. J. Climate,
13, 4018–4027.

——, 2004: An index of interannual precipitation variability in the
core of the North American monsoon region. J. Climate, 17,
4473–4480.

——, and J. W. Preston, 1997: Evidence for a relationship be-
tween spring snow cover and summer rainfall in New Mexico.
Geophys. Res. Lett., 24, 2207–2210.

——, and Coauthors, 2004: North American monsoon Model As-
sesment Project (NAMAP). NCEP/Climate Prediction Cen-
ter ATLAS 11, 32 pp. [Available from Climate Prediction
Center, 5200 Auth Road, Room 605, Camp Springs, MD 20746
and online at http://www.cpc.ncep.noaa.gov/research_papers/
ncep_cpc_atlas/11/atlas11.htm.]

Hastenrath, S., and L. Greischar, 1993: Further work on the pre-
diction of northeast Brazil rainfall anomalies. J. Climate, 6,
743–758.

Herdies, D. L., A. Da Silva, M. A. Silva Dias, and R. Nieto-
Ferreira, 2002: Moisture budget of the bimodal pattern of the
summer circulation over South America. J. Geophys. Res.,
107, 8075, doi:10.1029/2001JD000997.

Higgins, R. W., and W. Shi, 2000: Dominant factors responsible
for interannual variability of the summer monsoon in the
southwestern United States. J. Climate, 13, 759–776.

——, and ——, 2001: Intercomparison of the principal modes of
interannual and intraseasonal variability of the North Ameri-
can Monsoon System. J. Climate, 14, 403–417.

——, and NAME Science Working Group, 2003: The North
American Monsoon Experiment (NAME). Science and
Implementation Plan for NAME, 91 pp. [Available from Cli-
mate Prediction Center, 5200 Auth Road, Camp Springs, MD
20746 or online at http://www.joss.ucar.edu/name.]

——, Y. Yao, and X.-L. Wang, 1997: Influence of the North
American monsoon system on the U.S. summer precipitation
regime. J. Climate, 10, 2600–2622.

——, K. C. Mo, and Y. Yao, 1998: Interannual variability of the
U.S. summer precipitation regime with emphasis on the
southwestern monsoon. J. Climate, 11, 2582–2606.

——, Y. Chen, and A. V. Douglas, 1999: Interannual variability of
the North American warm season precipitation regime. J.
Climate, 12, 653–680.

——, A. Leetmaa, Y. Xue, and A. Barnston, 2000: Dominant
factors influencing the seasonal predictability of U.S. precipi-
tation and surface air temperature. J. Climate, 13, 3994–4017.

——, and Coauthors, 2003: Progress in Pan American CLIVAR
research: The North American Monsoon System. Atmosfera,
16, 29–65.

——, W. Shi, and C. Hain, 2004: Relationships between Gulf of

California moisture surges and precipitation in the south-
western United States. J. Climate, 17, 2983–2997.

Horel, J. D., A. N. Hahmann, and J. E. Geisler, 1989: An investi-
gation of the annual cycle of convective activity over tropical
Americas. J. Climate, 2, 1388–1403.

Hu, Q., and S. Feng, 2001: Climatic role of the southerly flow from
the Gulf of Mexico in interannual variations in summer rain-
fall in the central United States. J. Climate, 14, 3156–3170.

——, and ——, 2002: Interannual rainfall variations in the North
American summer monsoon region: 1900–98. J. Climate, 15,
1189–1202.

Huang, Z., and C. A. Lai, 1998: Multidecadal variability on the
southwestern United States monsoon precipitation in the
past century. Preprints, Ninth Conf. on the Interaction of the
Sea and Atmosphere, Phoenix, AZ, Amer. Meteor. Soc., 159–
166.

Huntingford, C., P. Harris, N. Gedney, P. Cox, R. Betts, J.
Marengo, and J. Gash, 2004: Investigating the potential of
Amazon “die-back” in a future climate using a GCM ana-
logue model. Theor. Appl. Climatol., 78, 177–186.

Jones, C., and J. Schemm, 2000: The influence of intraseasonal
variations on medium- to extended-range weather forecasts
over South America. Mon. Wea. Rev., 128, 486–494.

——, and L. M. Carvalho, 2002: Active and break phases in the
South American monsoon system. J. Climate, 15, 905–914.

Kayano, M. T., 2003: Low-level high-frequency modes in the
Tropical Atlantic and their relation to precipitation in the
equatorial South America. Meteor. Atmos. Phys., 83, 263–
276.

Kodama, Y.-M., 1992: Large-scale common features of sub-
tropical precipitation zones (the Baiu Frontal Zone, the
SPCZ, and the SACZ). Part I: Characteristics of subtropical
frontal zones. J. Meteor. Soc. Japan, 70, 813–835.

Koster, R., M. J. Suarez, and M. Heiser, 2000: Variance and pre-
dictability of precipitation at seasonal-to-interannual time-
scales. J. Hydrometeor., 1, 26–46.

Kousky, V. E., 1988: Pentad outgoing longwave radiation clima-
tology for the South American sector. Rev. Bras. Meteor., 3,
217–231.

Larson, J., Y. Zhou, and R. W. Higgins, 2005: Characteristics of
landfalling tropical cyclones in the United States and Mexico:
Climatology and interannual variability. J. Climate, 18, 1247–
1262.

Liebmann, B., and J. A. Marengo, 2001: The seasonality and in-
terannual variability of rainfall in the Brazilian Amazon ba-
sin. J. Climate, 14, 4308–4318.

——, H. Hendon, and J. Glick, 1994: The relationship between
tropical cyclones of the western Pacific and Indian Oceans
and the Madden-Julian Oscillation. J. Meteor. Soc. Japan, 72,
401–411.

——, G. N. Kiladis, J. A. Marengo, T. Ambrizzi, and J. D. Glick,
1999: Submonthly convective variability over South America
and the South Atlantic convergence zone. J. Climate, 12,
1877–1891.

——, ——, C. S. Vera, and A. C. Saulo, 2004a: Subseasonal varia-
tions of rainfall in South America in the vicinity of the low-
level jet east of the Andes and comparison to those in the
South Atlantic convergence zone. J. Climate, 17, 3829–3842.

——, and Coauthors, 2004b: An observed trend in central South
American precipitation. J. Climate, 17, 4357–4367.

15 OCTOBER 2006 V E R A E T A L . 4997



Madden, R. A., and P. R. Julian, 1994: Observations of the 40–
50-day tropical oscillation—A review. Mon. Wea. Rev., 122,
814–837.

Maddox, R. A., D. M. McCollum, and K. W. Howard, 1995:
Large-scale patterns associated with severe summertime
thunderstorms over central Arizona. Wea. Forecasting, 10,
763–778.

Magaña, V., J. A. Amador, and S. Medina, 1999: The midsummer
drought over Mexico and Central America. J. Climate, 12,
1577–1588.

Maloney, E. D., and D. L. Hartmann, 2000: Modulation of eastern
North Pacific hurricanes by the Madden–Julian oscillation. J.
Climate, 13, 1451–1460.

Mantua, J., S. Hare, Y. Zhang, J. M. Wallace, and R. Francis,
1997: A Pacific interdecadal climate oscillation with impacts
on salmon production. Bull Amer. Meteor. Soc., 78, 1069–
1080.

Mapes, B. E., T. T. Warner, and M. Xu, 2003a: Diurnal patterns of
rainfall in northwestern South America. Part III: Diurnal
gravity waves and nocturnal convection offshore. Mon. Wea.
Rev., 131, 830–844.

——, ——, ——, and A. J. Negri, 2003b: Diurnal patterns of rain-
fall in northwestern South America. Part I: Observations and
context. Mon. Wea. Rev., 131, 799–812.

Marengo, J. A., 1992: Interannual variability of surface climate in
the Amazon basin. Int. J. Climatol., 12, 853–863.

——, 2004: Interdecadal and long term-variations of rainfall in the
Amazon basin. Theor. Appl. Climatol., 78, 79–96.

——, and C. A. Nobre, 2001: The hydroclimatological framework
in Amazonia. Biogeochemistry of Amazonia, J. Richey, M.
McClaine, and R. Victoria, Eds., Springer-Verlag, 17–42.

——, B. Liebmann, V. Kousky, N. Filizola, and I. Wainer, 2001:
On the onset and end of the rainy season in the Brazilian
Amazon basin. J. Climate, 14, 833–852.

——, M. Douglas, and P. Silva Dias, 2002: The South American
low-level jet east of the Andes during the 1999 LBA-TRMM
and LBA-WET AMC campaign. J. Geophys Res., 107, 8079,
doi:10.1029/2001JD001188.

——, and Coauthors, 2003: Ensemble simulation of regional rain-
fall features in the CPTEC/COLA atmospheric GCM. Skill
and Predictability assessment and applications to climate pre-
dictions. Climate Dyn., 21, 459–475.

——, W. Soares, C. Saulo, and M. Nicolini, 2004: Climatology of
the low-level jet east of the Andes as derived from the
NCEP–NCAR reanalyses: Characteristics and temporal vari-
ability. J. Climate, 17, 2261–2280.

——, G. Fisch, I. Vendrame, I. Cervantes, and C. Morales, 2005:
On the diurnal and day-to-day variability of rainfall in South-
west Amazonia during the LBA-TRMM and LBA-WET
AMC Campaigns of summer 1999. Acta Amazonica, 34, 593–
603.

Mechoso, C. R., S. Lyons, and J. Spahr, 1990: The impact of sea
surface temperature anomalies on the rainfall in northeast
Brazil. J. Climate, 3, 812–826.

McCollum, D. M., R. A. Maddox, and R. A. Howard, 1995: Case
study of a severe mesoscale convective system in central Ari-
zona. Wea. Forecasting, 10, 643–665.

Mitchell, D. L., D. Ivanova, R. Rabin, T. J. Brown, and K. Red-
mond, 2002: Gulf of California sea surface temperatures and
the North American monsoon: Mechanistic implications
from observations. J. Climate, 15, 2261–2281.

Mo, K. C., 2001: Adaptive filtering and prediction of intraseasonal
oscillations. Mon. Wea. Rev., 129, 802–817.

Mock, C. J., 1996: Climatic controls and spatial variations of pre-
cipitation in the western United States. J. Climate, 9, 1111–
1125.

Molinari, J., and D. Vollaro, 2000: Planetary- and synoptic-scale
influences on eastern Pacific tropical cyclogenesis. Mon. Wea.
Rev., 128, 3296–3307.

——, D. Knight, M. Dickinson, D. Vollaro, and S. Skubis, 1997:
Potential vorticity, easterly waves, and eastern Pacific tropi-
cal cyclogenesis. Mon. Wea. Rev., 125, 2699–2708.

——, D. Vollaro, S. Skubis, and M. Dickinson, 2000: Origins and
mechanisms of eastern Pacific tropical cyclogenesis: A case
study. Mon. Wea. Rev., 128, 125–139.

Moura, A. D., and E. J. Shukla, 1981: On the dynamics of the
droughts in northeast Brazil: Observations, theory and nu-
merical experiments with a general circulation model. J. At-
mos. Sci., 38, 2653–2673.

Mozer, J. B., and J. A. Zehnder, 1996: Lee vorticity production by
large-scale tropical mountain ranges. Part I: Eastern North
Pacific tropical cyclogenesis. J. Atmos. Sci., 53, 521–538.

Negri, A. N., R. F. Adler, E. J. Nelkin, and G. J. Huffman, 1994:
Regional rainfall climatology derived from Special Sensor
Microwave Imager (SSM/I) data. Bull. Amer. Meteor. Soc.,
75, 1165–1182.

Nesbitt, S. W., and E. J. Zipser, 2003: The diurnal cycle of rainfall
and convective intensity to three years of TRMM measure-
ments. J. Climate, 16, 1456–1475.

Nicolini, M., P. Salio, G. Ulke, J. Marengo, M. Douglas, J. Paegle,
and E. Zipser, 2004: South American low-level jet diurnal
cycle and three-dimensional structure. CLIVAR Exchanges,
Vol. 9, No. 1, International CLIVAR Project Office,
Southampton, United Kingdom, 6–9.

Nieto-Ferreira, R. N., and W. H. Schubert, 1997: Barotropic as-
pects of ITCZ breakdown. J. Atmos. Sci., 54, 261–285.

——, T. M. Rickenbach, D. L. Herdies, and L. M. V. Carvalho,
2003: Variability of South American convective cloud sys-
tems and tropospheric circulation during January–March
1998 and 1999. Mon. Wea. Rev., 131, 961–973.

Nogués-Paegle, J., and K.-C. Mo, 1997: Alternating wet and dry
conditions over South America during summer. Mon. Wea.
Rev., 125, 279–291.

——, and ——, 2002: Linkages between summer rainfall variabil-
ity over South America and sea surface temperature anoma-
lies. J. Climate, 15, 1389–1407.

——, ——, and J. Paegle, 1998: Predictability of the NCEP–
NCAR reanalysis model during austral summer. Mon. Wea.
Rev., 126, 3135–3152.

——, L. A. Byerle, and K. C. Mo, 2000: Intraseasonal modulation
of South American summer precipitation. Mon. Wea. Rev.,
128, 837–850.

——, and Coauthors, 2002: Progress in Pan American CLIVAR
research: Understanding the South American Monsoon. Me-
teorologica, 27, 3-30.

Okabe, I. T., 1995: The North American Monsoon. Ph.D. disser-
tation, University of British Columbia, 146 pp.

Paegle, J., and Coauthors, 2004: Modeling studies related to
SALLJEX. CLIVAR Exchanges, Vol. 9, No. 1, International
CLIVAR Project Office, Southampton, United Kingdom,
20–23.

4998 J O U R N A L O F C L I M A T E — S P E C I A L S E C T I O N VOLUME 19



Petersen, W. A., S. Nesbitt, R. J. Blakeslee, R. Cifelli, P. Hein,
and S. A. Rutledge, 2002: TRMM observations of intrasea-
sonal variability in convective regimes over the Amazon. J.
Climate, 15, 1278–1293.

Pisciottano, G., A. Diaz, G. Cazes, and C. R. Mechoso, 1994: El
Niño–Southern Oscillation impact on rainfall in Uruguay. J.
Climate, 7, 1286–1302.

Rao, V. B., and K. Hada, 1990: Characteristics of rainfall over
Brazil. Annual variations and connections with the Southern
Oscillation. Theor. Appl. Climatol., 42, 81–91.

Robertson, A., and C. R. Mechoso, 2003: Circulation regimes and
low-frequency oscillations in the South Pacific dector. Mon.
Wea. Rev., 131, 1566–1576.

——, ——, and Y.-J. Kim, 2000: The influence of Atlantic SST
anomalies on the North Atlantic Oscillation. J. Climate, 13,
122–138.

Rodwell, M. J., and B. J. Hoskins, 2001: Subtropical anticyclones
and summer monsoons. J. Climate, 14, 3192–3211.

Ropelewski, C. F., and M. S. Halpert, 1986: North American pre-
cipitation and temperature patterns associated with the El
Nino/Southern Oscillation (ENSO). Mon. Wea. Rev., 114,
2352–2362.

Rowson, D. R., and S. J. Colucci, 1992: Synoptic climatology of
thermal low-pressure systems over South-Western North
America. J. Climatol., 12, 529–545.

Salio, P., M. Nicolini, and A. C. Saulo, 2002: Chaco low-level jet
events characterization during the austral summer season. J.
Geophys. Res., 107, 4816, doi:10.1029/2001JD001315.

Saulo, A. C., M. Nicolini, and S. C. Chou, 2000: Model character-
ization of the South American low-level flow during the
1997–1998 spring-summer season. Climate Dyn., 16, 867–881.

——, L. J. Ferreira, J. Mejia, and M. Seluchi, 2004: Description of
the Thermal Low characteristics using SALLJEX special ob-
servations. CLIVAR Exchanges, Vol. 9, No. 1, International
CLIVAR Project Office, Southampton, United Kingdom,
9–11.

Schmitz, J. T., and S. L. Mullen, 1996: Water vapor transport as-
sociated with the summertime North American monsoon as
depicted by ECMWF analyses. J. Climate, 9, 1621–1634.

Seluchi, M., and A. C. Saulo, 1998: Possible mechanisms yielding
an explosive coastal cyclogenesis over South America: Ex-
periments using a limited area model. Aust. Meteor. Mag., 47,
309–320.

——, E. Marcelo, A. Saulo, C. M. Nicolini, and P. Satyamurty,
2003: The northwestern Argentinean low: A study of two
typical events. Mon. Wea. Rev., 131, 2361–2378.

Silva Dias, M. A. F., and Coauthors, 2002: Clouds and rain pro-
cesses in a biosphere-atmosphere interaction context in the
Amazon Region. J. Geophys. Res., 107, 8072, doi:10.1029/
2001JD000335.

Silva Dias, P. L., J. P. Bonatti and V. E. Kousky, l987: Diurnally
forced tropical tropospheric circulation over South America.
Mon. Wea. Rev., 115, 1465-1478.

Silvestri, G. E., and C. S. Vera, 2003: Antarctic Oscillation signal
on precipitation anomalies over southeastern South America.
Geophys. Res. Lett., 30, 2115, doi:10.1029/2003GL018277.

Sorooshian, S., X. Gao, K. Hsu, R. A. Maddox, Y. Hong, H. V.
Gupta, and B. Imam, 2002: Diurnal variability of tropical
rainfall retrieved from combined GOES and TRMM satellite
information. J. Climate, 15, 983–1001.

Stensrud, D. J., R. L. Gall, S. L. Mullen, and K. W. Howard, 1995:
Model climatology of the Mexican monsoon. J. Climate, 8,
1775–1794.

——, ——, and M. K. Nordquist, 1997: Surges over the Gulf of
California during the Mexican monsoon. Mon. Wea. Rev.,
125, 417–437.

Sugahara, S., 1991: Flutuações interanuais, sazonais e intrasa-
zonais de precipitação no estado de São Paulo. Ph.D thesis,
Universidade de São Paulo, 140 pp.

Uvo, C. R. B., C. A. Repelli, S. Zebiak, and Y. Kushnir, 1998: The
relationship between tropical Pacific and Atlantic SST and
northeast Brazil monthly precipitation. J. Climate, 11, 551–
562.

Velasco, I., and J. M. Fritsch, 1987: Mesoscale convective com-
plexes in the Americas. J. Geophys. Res., 92, 9591–9613.

Vera, C., 2004: Introduction to the South American Low Level Jet
Experiment (SALLJEX). CLIVAR Exchanges, Vol. 9, No. 1,
International CLIVAR Project Office, Southampton, United
Kingdom, 3–4.

——, and C. Nobre, 1999: On the dynamics associated with the
onset of the wet season in Central and Southeastern Brazil.
2nd Session of the CLIVAR/VAMOS Panel, WCRP Rep.
9/99, ICPO Rep. 30, A41–A43.

——, and P. K. Vigliarolo, 2000: Variations of South America
summer circulation on subseasonal time scales. Preprints,
Sixth Conf. on Southern Hemisphere Meteorology and Ocean-
ography, Santiago, Chile, Amer. Meteor. Soc., 170–171.

——, ——, and E. H. Berbery, 2002: Cold season synoptic-scale
waves over subtropical South America. Mon. Wea. Rev., 130,
684–699.

——, G. Silvestri, V. Barros, and A. Carril, 2004: Differences in El
Niño response over the Southern Hemisphere. J. Climate, 17,
1741–1753.

Viramontes, D., and L. Descroix, 2003: Changes in the surface
water hydrologic characteristics of an endoreic basin of
northern Mexico from 1970 to 1998. Hydrol. Processes, 17,
1291–1306.

Wagner, R., 1996: Decadal-scale trends in mechanisms controlling
meridional sea surface temperature gradients in the tropical
Atlantic. J. Geophys. Res., 101, 16 683–16 694.

Wang, H., and R. Fu, 2002: Cross-equatorial flow and seasonal
cycle of precipitation over South America. J. Climate, 15,
1591–1608.

Wheeler, M., and K. M. Weickmann, 2001: Real-time monitoring
and prediction of modes of coherent synoptic to intraseasonal
tropical variability. Mon. Wea. Rev., 129, 2677–2694.

Xie, P., and P. A. Arkin, 1997: Global precipitation: A 17-year
monthly analysis based on gauge observations, satellite esti-
mates, and numerical model outputs. Bull. Amer. Meteor.
Soc., 78, 2539–2558.

Yang, S., K. M. Lau, and M. Sankar-Rao, 1996: Precursory signals
associated with the interannual variability of the Asian sum-
mer monsoon. J. Climate, 9, 949–969.

Yu, B., and J. M. Wallace, 2000: The principal mode of interan-
nual variability of the North American Monsoon System. J.
Climate, 13, 2794–2800.

Yu, J.-Y., and C. R. Mechoso, 1999: Links between annual varia-
tions of Peruvian stratocumulus clouds and of SST in the
eastern equatorial Pacific. J. Climate, 12, 3305–3318.

Zehnder, J., 1991: The interaction of planetary-scale tropical east-

15 OCTOBER 2006 V E R A E T A L . 4999



erly waves with topography: A mechanism for the initiation
of tropical cyclones. J. Atmos. Sci., 48, 1217–1230.

——, D. M. Powell, and D. L. Ropp, 1999: The interaction of
easterly waves, orography, and the intertropical convergence
zone in the genesis of eastern Pacific tropical cyclones. Mon.
Wea. Rev., 127, 1566–1585.

Zhou, J., and W. K.-M. Lau, 1998: Does a monsoon climate exist
over South America? J. Climate, 11, 1020–1040.

——, and ——, 2001: Principal modes of interannual and decadal
variability of summer rainfall over South America. Int. J.
Climatol., 21, 1623–1644.

Zipser, E., P. Salio, and M. Nicolini, 2004: Mesoscale convective
systems activity during SALLJEX and the relationship with
SALLJ events. CLIVAR Exchanges, Vol. 9, No. 1, Interna-
tional CLIVAR Project Office, Southampton, United King-
dom, 14–18.

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