MAY 2004
RICKENBACH
1201
Nocturnal Cloud Systems and the Diurnal Variation of Clouds and Rainfall in
Southwestern Amazonia
THOMAS M. RICKENBACH
Joint Center for Earth Systems Technology, University of Maryland, Baltimore County, Baltimore, and Mesoscale Processes Branch,
NASA GSFC, Greenbelt, Maryland
(Manuscript received 23 July 2003, in final form 19 November 2003)
ABSTRACT
This paper examines the origins of a secondary nocturnal maximum in cloudiness and precipitation in southwestern Amazonia, a diurnal feature observed previously by many investigators. Analysis is based on satellite,
radar, sounding, and profiler observations of precipitating systems and cloudiness from the Tropical Rainfall
Measuring Mission Large-Scale Biosphere–Atmosphere (TRMM-LBA) and the coincident Wet-Season Atmospheric Mesoscale Campaign (WETAMC) field programs during the early 1999 wet season. The general finding
is that following the collapse of the nearly ubiquitous and locally generated afternoon (‘‘noon balloon’’) convection, organized deep convection contributes to a postmidnight maximum in raining area and high cloudiness,
and to a lesser extent rainfall. Nocturnal convective systems have the effect of weakening and delaying the
onset of the following afternoon’s convection. Many of these nocturnal convective events are traced to largescale squall lines, which propagate westward thousands of kilometers from their point of origin along the northeast
coast of Brazil. In addition, a previously undescribed nocturnal stratiform drizzle phenomenon, generated above
the melting layer independently from deep convection, contributes significantly to nocturnal cloud cover. Results
from this study underscore the complex influence of propagating large-scale organized convection in locally
modulating the diurnal variation in clouds and rain. The greatest significance of the nocturnal drizzle may be
the potential effect on the diurnal radiation budget by the extensive midlevel nocturnal clouds rather than their
marginal contribution to nocturnal rainfall.
1. Introduction
The diurnal variation of clouds and precipitation over
tropical landmasses is the net expression of a rich variety
of phenomena spanning many time and space scales.
Solar heating of the land surface increases sensible heat
flux and thermodynamic instability, leading to vertical
mixing of water vapor in the deepening boundary layer
during the morning and favoring afternoon precipitating
convection (Wallace 1975; Betts 1976). Dynamic processes, such as land/water circulations, density currents,
and orographic lifting, are crucial to actually initiating
moist convection by lifting parcels to their level of positive buoyancy (Wallace 1975). Preexisting convective
systems may prime the atmosphere for further convection by lowering convective inhibition due to the net
upward displacement resulting from the atmosphere’s
response to heating from nearby convective systems.
Thus convective systems increase the chance that surface heating and local circulations will succeed at initiating more convection (Mapes 1993). Once formed,
subsequent upscale growth and mesoscale organization
Corresponding author address: Dr. Thomas M. Rickenbach, Code
912, NASA Goddard Space Flight Center, Greenbelt, MD 20771.
E-mail: [email protected]
q 2004 American Meteorological Society
strongly influences the timing of maximum rainfall and
cloudiness (McAnelly and Cotton 1989).
Over the last decade, advances in satellite remote
sensing and increased deployment of field instrumentation have led to a detailed view of diurnal changes in
rainfall and cloudiness. Satellite studies and surface observations have provided guidance for model parameterization of convective processes (Lin et al. 2000; Yang
and Slingo 2001; Betts and Jakob 2002). There has been
a recent focus on convective cloud systems in the Amazon basin, in part because of field efforts to study the
interaction between convection, land processes, topography, and large-scale circulations (Baidya Roy and Avissar 2002; Rickenbach et al. 2002; Laurent et al. 2002).
The Tropical Rainfall Measuring Mission Large-Scale
Biosphere–Atmosphere (TRMM-LBA) and the coincident Wet-Season Atmospheric Mesoscale Campaign
(WETAMC) field programs during the 1999 wet season
(Silva Dias et al. 2002) provided a detailed view of the
structure and environment of convective systems in
southwestern Amazonia.
Many studies of the diurnal cycle of precipitation and
cloudiness over the Amazon indicated a secondary nocturnal maximum in addition to the primary afternoon
maximum. A composite of two wet seasons of TRMM
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rain estimates from the precipitation radar (PR) over the
Amazon (Lin et al. 2000) indicated that rainfall increased at 2300 LT to a broad nocturnal maximum,
though weaker than the afternoon maximum following
the explosive increase between 1100 LT and local noon.
Lin et al. (2000) postulated that the weaker nocturnal
maximum was associated with stratiform rain enhanced
through instability driven by nocturnal radiative cooling
of the extensive anvil cloud tops. Negri et al. (2002)
found a similar nocturnal maximum in rainfall using a
TRMM-adjusted geostationary infrared rainfall estimate
applied to the Amazon basin. They hypothesized that
the nocturnal peak was associated with an enhancement
of rainfall along the wide Amazon River via a land/river
circulation favoring nocturnal convection over the river.
Nesbitt and Zipser (2003) used TRMM PR and passive
microwave measurements to partition rainfall into contributions from isolated convection with little ice-scattering signature and mesoscale convective systems
(MCSs) with extensive thick anvil (ice) cloud. In the
composite of all tropical land regions, isolated convection with and without ice aloft had peak rainfall in the
afternoon, while rainfall from MCSs was greatest
around midnight, presumably from the upscale growth
of afternoon convection.
A further complication to this picture was that the
phase of the diurnal cycle of clouds and rainfall varied
across Amazonia. Satellite retrievals of the diurnal cycle
of precipitation and cloudiness over the Amazon basin
revealed an alternating banded structure in the diurnal
phase, from the northeast coast to western Amazonia.
Negri et al. (1994) used passive microwave measurements to construct day versus evening precipitation
composites for three wet seasons. They found that the
prevalence of afternoon rainfall was not uniform over
the basin, but instead was focused in parallel bands
along the entire northeast coast of Brazil, central Amazonia, and western Amazonia. Wet-season composites
of the diurnal phase from infrared cloud-top temperature
data showed a similar pattern (Garreaud and Wallace
1997; Yang and Slingo 2001). Garreaud and Wallace
(1997) posited that the banded structure in diurnal phase
might be influenced by gravity wave propagation from
regions of strong diurnal modulation on both sides of
the basin (orographic convection in the Andes to the
west, and afternoon sea-breeze convection along the
northeast coast to the east).
Other investigations further suggested that the phase
propagation of the diurnal rainfall maximum in Amazonia was tied to the effects of long-lived, propagating
convective systems. Greco et al. (1990) found that in
central Amazonia 1200 km from the northeast coast,
convective systems that formed locally produced a rainfall maximum around 1800 LT, while rainfall associated
with propagating squall-line systems that formed along
the northeast coast produced a maximum between 1000
and 1400 LT. They hypothesized that the late morning
maximum corresponded to the time of arrival of the
VOLUME 132
coastal squall lines to central Amazonia. Cutrim et al.
(2000) analyzed a different year of rain gauge data to
show that a station in central Amazonia had a maximum
near local noon, much earlier than the 1600 LT maximum at a station on the northeast coast, possibly tied
to the timing of coastal squall-line passage as described
in Greco et al. (1990). A satellite climatology of these
large-scale westward propagating coastal systems presented by Cohen et al. (1995) showed that though they
were more frequent (or more easily identified) during
the dry season, they occurred all year round and could
be tracked at times more than 2000 km inland (westward). The results of Garreaud and Wallace (1997) suggested that as large-scale coastal squall lines propagate
westward to the basin interior, the diurnal maximum in
cloudiness follows them in space and time, both during
the wet and dry seasons. Similar diurnal rainfall phase
propagation occurred across West Africa for related reasons (Shinoda et al. 1999; Fortune 1980).
Recent results from the TRMM-LBA campaign in
southwestern Amazonia provided additional insight. Diurnal composites of radar-derived rain area and conditional rainfall rate were produced for periods of northwesterly low-level flow in Rondônia State, Brazil [when
the South Atlantic convergence zone (SACZ) penetrated
the deep Tropics] and were compared with the interim
period (Rickenbach et al. 2002). In addition to the afternoon maxima in rainfall and rain area, they found for
both regimes another large maximum in rain area after
local midnight following a minimum near 2300 LT. The
presence of this minimum suggested nocturnal resurgence of light rainfall, well after the decay of the afternoon systems, though the reasons for this were not
clear. Machado et al. (2002) also noted a similar secondary nocturnal peak in radar rain area and total cloud
cover during the TRMM-LBA campaign.
These studies highlighted several unresolved questions concerning the nocturnal maximum in clouds and
precipitation over tropical South America. What was
the origin of the secondary nocturnal maximum in rainfall and cloud area? How important were stratiform precipitation processes in the nocturnal maximum of clouds
and rainfall? Did propagating coastal squall lines influence nocturnal rainfall in southwestern Amazonia? The
goal of this paper is to investigate nocturnal precipitation in Amazonia with those specific questions in mind.
This study takes the approach of composite and case
study analysis of TRMM-LBA surface observations and
geostationary satellite data for nearly 2 months in southwestern Amazonia during the early 1999 wet season.
Results from this study highlight the complex influence
of propagating large-scale organized convection in locally modulating the diurnal variation in clouds and rain.
These findings may help to guide the interpretation of
multiyear diurnal rainfall and cloudiness composites
from satellite platforms such as TRMM and may lead
to improvements in the representation of precipitation
processes in model parameterizations of convection.
MAY 2004
RICKENBACH
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2. Data
The TRMM-LBA/WETAMC field campaign was
conducted in Rondônia in southwestern Amazonia during January and February 1999. The main objectives of
these joint experiments were to examine biosphere–atmosphere coupling and to provide validation for TRMM
algorithms in a tropical continental environment (for
details see Silva Dias et al. 2002). In this study, volumes
of radar reflectivity data, taken approximately every 10
min from the National Aeronautics and Space Administration (NASA)C-band scanning radar [known as the
Tropical Ocean Global Atmosphere (TOGA) radar],
were used to examine the three-dimensional structure
of convection between 11 January and 28 February.
Radar reflectivity data were quality controlled to remove
anomalous echo, and calibration biases were corrected
by comparison with the TRMM precipitation radar
(Rickenbach et al. 2002; Anagnostou and Morales
2002). The data were interpolated to a 1 km 3 1 km
3 1 km grid for analysis. Spatial gradients and local
maxima provided the basis to partition the reflectivity
field into regions of active convective cells and weaker,
horizontally uniform stratiform rain (see Rickenbach et
al. 2002 for details). Analysis was restricted to within
a range of 130 km from the radar since beam broadening
and gaps between radar scans degraded the usefulness
of the data beyond that range. Reflectivity data at a
height of 2 km above ground level (AGL) were converted to rainfall rate using an empirical relationship
based on surface disdrometer data (Tokay et al. 2002).
Only relative differences rather than absolute values of
rainfall rate are emphasized in this analysis. Rain area
was based on the areal coverage of reflectivity greater
than 10 dBZ (about 0.5 mm h 21 , or roughly speaking,
drizzle).
Infrared brightness temperature data for the Amazon
basin (every 30 min, nominally 5-km spacing) from the
Geostationary Operational Environmental Satellite-8
(GOES-8) revealed the height, organization, and motion
of cloud tops associated with convective systems and
other cloud features. These data were obtained from the
NASA Goddard Space Flight Center’s Distributed Active Archive Center (GSFC DAAC) and navigated to
earth coordinates in order to produce imagery and regional composites over different areas of interest. Composites were made within the radar coverage area and
within a 108 diameter circle covering western and central
Amazonia inclusive of the radar area. Sounding data
from the Abracos site (collocated with the radar) and
the Rancho Grande site (at Ariquemes, 70 km northwest
of the radar) provided thermodynamic and wind profiles
for two case studies presented in this work [see Halverson et al. (2002) for details on the radiosonde network in the TRMM-LBA campaign]. Figure 1 shows
the location of the observational network and analysis
regions used in this study.
FIG. 1. (top) Map of the Amazon basin with topographic contours
(every 100 m), latitude/longitude grid (every 58), and several geographic regions discussed in the text. The small circle (150-km radius)
shows the TOGA radar areal coverage, within which composites of
rainfall and IR pixel counts were made. Infrared pixel count composites were also computed in the large (108 diameter) circle for
comparison. (bottom) The location of the radar, sounding launches,
and profiler observations within the TOGA area (0.58 latitude–longitude grid).
3. Results
We begin by presenting general features of the diurnal
cycle of rainfall and cloudiness, from 11 January 1999
to 28 February 1999 (49 days). Diurnal composites of
conditional rainfall rate (rainfall within the raining area),
rain area, and vertical reflectivity structure (Fig. 2) were
constructed from the radar reflectivity volumes and
compared to composites of cloudiness from geostationary satellite. Rain area, rainfall rate, and mean echo-top
height were separated into convective and stratiform
components. Included in the radar parameters was the
mean height of the 30-dBZ surface, which is an indicator
of the vertical intensity of convective cells (e.g., DeMott
and Rutledge 1998). A greater height of the 30-dBZ
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FIG. 3. (a) Infrared pixel count (proportional to area) within the
TOGA radar area for all days at thresholds of 250 (solid), 230
(dotted), and 210 K (dashed). (b) Infrared pixel count (proportional
to area) within a 108 diameter circle (cf. Fig. 1) for all days at thresholds of 250 (solid), 230 (dotted), 210 K (dashed).
surface suggests larger concentrations and/or larger particle sizes of condensate in the mixed-phase region of
the cell, and thus stronger convective updrafts. Infrared
cloudiness area (number of pixels) was calculated for
various brightness temperature thresholds. The colder
thresholds represented convective cloud tops and cold
anvil cloud, while the warmest thresholds were related
to the total cloudiness including midlevel clouds.
a. Radar and infrared diurnal composites for the
TRMM-LBA campaign (49 days)
FIG. 2. Diurnal composites from the TOGA radar for the entire
experiment. (a) Fraction of radar area covered by rain (dBZ . 10 at
2 km AGL): convective is dotted line, stratiform is dashed, and total
is solid. (b) Conditional rainfall rate (mm h 21 ): convective is dotted
line, stratiform is dashed, and total is solid. (c) Echo-top height:
convective is dotted line, stratiform is dashed, and 30-dBZ height is
solid.
The radar observations showed that, following a midmorning minimum, a sharp increase in total rain area
(Fig. 2a) between 1000 and 1400 LT was coincident
with a similarly sharp increase in convective rainfall
(Fig. 2b), echo-top height, and 30-dBZ surface height
(Fig. 2c) over the same time period. This was the signature of rapid, widespread formation of precipitating
convective cells near local noon (hereafter ‘‘noon bal-
MAY 2004
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RICKENBACH
TABLE 1. Local hour of the peak in midlevel cloudiness (250-K IR threshold) and mean 12-h (2200–1000LT) radar-derived rain rate for
the 25 days of nocturnal systems within the radar area. Each event is categorized as nocturnal convection (bold font) or nocturnal drizzle.
The region of origin of large-scale squall lines is indicated. The two italicized events are presented in detail in section 3b.
Nocturnal
convection
12
15
18
19
22
23
24
25
26
29
30
1
2
4
7
8
11
15
17
18
19
23
24
27
28
Jan
Jan
Jan
Jan
Jan
Jan
Jan
Jan
Jan
Jan
Jan
Feb
Feb
Feb
Feb
Feb
Feb
Feb
Feb
Feb
Feb
Feb
Feb
Feb
Feb
Nocturnal
drizzle
0100
0000
0500
0400
0300
0300
0000
0600
0300
0600
0200
0300
0000
0300
0500
0200
0300
0200
0100
0400
0000
0400
0600
0700
0400
Mean rain rate
2200–1000 LT
(mm h21 )
0.0016
0.23
0.25
0.030
0.27
0.024
0.15
0.41
0.21
0.039
0.27
0.019
0.14
0.098
0.41
0.20
0.018
0.49
0.011
1.05
0.056
0.046
0.53
0.49
0.061
loon,’’ a term coined by E. Williams). Convective rain
area and rainfall decreased after 1600 LT, while stratiform rain area remained constant and mean echo-top
height continued to increase. These observations suggested that by 1600 LT, active convection weakened,
stratiform anvil clouds expanded in area, and rainfall
decreased, all pointing to the decay stage of the noonballoon convection. By about 2200 LT, the composites
suggested that the afternoon convective activity had
come to an end, having collapsed in vertical intensity
and diminished in area and rainfall. Animation of radar
imagery for each day showed a consistency with this
general scenario for most days.
After 2200 LT, the radar composites indicated a resurgence of activity. Stratiform rain area increased
sharply again between 2300 and 0300 LT, corresponding
in time with increased mean echo-top height and a slight
rise in convective intensity (mean 30-dBZ surface
height). The 0300 LT stratiform rain area maximum was
equal in amplitude to the afternoon maximum at 1500
LT. However, the convective component of rain area
showed only a weak and gradual increase, in contrast
to the afternoon situation, with a 0300 LT maximum.
Rainfall rate showed a weak downward trend after 2200
LT into the early morning hours, with some fluctuations
in convective rain during that time, particularly near
0200 LT. By 0600 LT, this nocturnal resurgence had
diminished, ending in a general cessation of all activity
by 0900 LT. The scenario suggested by these observa-
Origin of large-scale
squall line
Midlevel cloud
over western Amazon?
yes
SE highlands
NE coast
NE coast
NE coast
yes
yes
NE coast
NE coast
yes
Western Amazon
tions was of weak convective regeneration and a large
expansion of stratiform rain area after midnight. Perusal
of radar imagery revealed that nocturnal convection was
indeed common, but in contrast to the afternoon convection there was no single or consistent scenario to
explain what was happening at night. This will be explored further in section 3b.
Diurnal composites of infrared pixel count (areal coverage) at various brightness temperature thresholds added insight into the nocturnal activity and provided an
independent perspective from the radar. The afternoon
noon-balloon convection left a clear cloud-top area signature in each IR threshold (Fig. 3a), with a maximum
in cloud cover between 1700 and 1800 LT. This was
consistent with the time of maximum mean radar echotop height, expected since radar echo-top patterns typically corresponded well to IR cloud-top structure within
convective systems (Zipser 1988; Rickenbach 1999).
The cloud-top area maximum occurred 2–3 h after the
maximum in low-level rain area and rainfall, as found
by the TRMM-LBA observations of Negri et al. (2002)
and Machado et al. (2002), and quite consistent with
the general evolution of developing, heavily raining
convective cells to extensive stratiform anvil cold cloud
shields (Zipser 1988; McAnelly and Cotton 1989; Yuter
and Houze 1998; Rickenbach 1999).
Afternoon cloudiness diminished during the early
evening hours; however, the cloud coverage at 250 K
(low and midlevel cloud) began to increase again by
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FIG. 4. GOES-8 infrared (channel 4) satellite image over the Amazon basin for (a) 1415 UTC (1015 LT at
the TOGA radar) 17 Feb and (b) 0215 UTC (2215 LT at the TOGA radar, 12 h later) 18 Feb. White circle
shows the TOGA radar area, and arrows indicate the squall line’s leading edge, as described in the text. Image
shows 58 latitude and longitude lines; color scale represents brightness temperature (K).
2200 LT, near the time of the stratiform rain area increase seen by the radar. The dominant nocturnal cloud
maximum was in the low and midlevel clouds (250 K)
at 0100 LT, which occurred 1 h before the peak in radar
echo-top height and 2 h before the rain area maximum.
Only later, at 0300 LT, a weak maximum was observed
at the colder cloud area threshold (210 K; Fig. 3a),
which matched the mean radar echo-top area maximum
and was suggestive of some deep convective activity.
In summary, the IR and radar composites were quite
consistent with one another and suggested a rapid expansion of midlevel stratiform cloud, with a relatively
MAY 2004
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FIG. 5. Time series (local time) of radar parameters for 18 Feb: (a) fraction of radar area covered by rain, as in Fig. 2a,
(b) conditional rain rate (mm h 21 ), as in Fig. 2b, (c) echo-top height and 30-dBZ surface height, as in Fig. 2c, and (d) IR
pixel count within the radar area. In (a)–(d), the solid line represents 250 K, dotted represents 230 K, and dashed represents
210 K.
weak convective resurgence after midnight. These observations led to the conclusion that the nocturnal cloud
systems were decoupled from the afternoon convection
and that the nighttime systems evolved differently than
the afternoon convective activity, with a weak convective maximum an hour or two after the expansion of
midlevel stratiform cloud. This evolution of the nocturnal systems appeared to be in contrast with our understanding of the development of mesoscale convective
systems, where a peak in convective rain followed by
the expansion of stratiform cloud would be expected
(Smull and Houze 1985; Rasmussen and Rutledge
1993). This anomaly was likely explained by the juxtaposition of two types of nocturnal precipitating systems, as will be discussed in section 3b.
One possible source of bias in these composites was
that the relatively small areal coverage of the radar in-
troduced a sampling bias from partially sampled or
propagating convective systems. For example, the
chance timing of convective systems entering the radar’s
view at arbitrary life cycle stages might account for the
nocturnal observations. One way to address this was to
reproduce the IR composites for a much larger area, to
ensure that the life cycle of one or several mesoscale
convective systems was included, yet not so large as to
include regions of potentially distinct diurnal rainfall
characteristics (such as the Andes Mountains or the
northeast coastal region). A 108 diameter circular region
centered at 108S, 608W, which included the radar area
(but was 50 times larger in area), was used for this
purpose, as shown in Fig. 1. Perusal of IR imagery
indicated that multiple mesoscale cloud shields are commonly observed simultaneously in this large area, so the
problem of cloud systems propagating in or out of the
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FIG. 7. Radar and satellite images at 0415 UTC (0015 LT) 18 Feb
(a) GOES-8 infrared satellite image, with TOGA radar area (circle),
Abracos sounding site (X), Rancho Grande sounding site (square),
and profiler location (triangle) (b) (top) Radar reflectivity map at 2
km AGL. Dashed line shows the location of the (bottom) vertical
cross section.
FIG. 6. Abracos sounding skew T diagrams for (a) 0300 UTC 18
Feb, prior to squall-line passage and (b) 1200 UTC 18 Feb, after
squall-line passage. Black lines represent the temperature and dewpoint traces, and the curved gray line shows the path of moist adiabatic ascent above the level of free convection. Values of CAPE (J
kg 21 ) and convective inhibition (CINE; J kg 21 ) are given below each
plot. Short and long wind barbs represent 5 and 10 m s 21 , respectively.
composite region was greatly diminished. The resulting
diurnal composite of IR cloud-top area (Fig. 3b) was
nearly identical to that of the IR composite within the
radar area. The 0100 LT maximum in midlevel cloud
area (250 K) and 0300 LT maxima at the colder thresholds were all present (though slightly weaker), as were
the afternoon maxima. This consistency between the two
IR diurnal composites gave confidence that system propagation or partially sampled life cycle did not contaminate the radar (and IR within the radar area) diurnal
composites.
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RICKENBACH
FIG. 8. As in Fig. 7, but for 0615 UTC (0215 LT) 18 Feb.
b. Nocturnal precipitation events
The next step toward understanding the nocturnal
rainfall and cloudiness maxima was to look for consistent patterns in the daily time series of radar and IR
parameters and animations of radar and IR imagery.
Examination of the midlevel IR cloud threshold (250
K) time series showed that 38 of the 49 experiment days
had a clear and distinct postmidnight maximum in
cloudiness. Of these 38 days, 13 had significant gaps
(at least 10 missing hours in a 24-h period) in either
the IR or radar data. These days were removed from
subsequent analysis, leaving 25 days in which the nocturnal cloudiness maxima were well sampled. Careful
perusal of radar and IR image animations and daily time
series suggested two general scenarios qualitatively de-
1209
FIG. 9. As in Fig. 7, but for 0815 UTC (0415 LT) 18 Feb.
scribing the nocturnal systems. On many of these nights,
one or more squall line systems propagated into the
radar view. These systems typically consisted of either
a large squall-line system, or a series of smaller groups
of active convection (hereafter referred to as ‘‘nocturnal
convection’’), with associated stratiform rain. On other
nights, large regions of light stratiform precipitation and
midlevel cloudiness formed within and well beyond the
radar area, with no clearly associated deep convection
(termed ‘‘nocturnal drizzle’’ in this paper). During the
TRMM-LBA campaign, many scientists in the field observed extensive nocturnal drizzle and midlevel overcast
that at times persisted after sunrise. Such events were
quite distinct from nocturnal convective systems, and
of sufficient frequency and uniqueness to warrant further
investigation.
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FIG. 10. As in Fig. 4, but for (a) 0145 UTC (2145 LT at the TOGA radar 18 Feb) 19 Feb and (b) 0945 UTC
(0545 LT at the TOGA radar) 19 Feb.
Table 1 presents a listing of the two general categories
of nocturnal systems (nocturnal convection and nocturnal drizzle) observed in the early morning hours for the
25 days. Maximum cloudiness occurred between 0000
and 0700 LT, with no specific time favored by either
category of nocturnal systems. The 12-h (2200–1000
LT) mean rainfall rate tended to be about 5 to 10 times
less for nocturnal drizzle events. Occurrence frequency
was nearly the same (14 nocturnal convection days versus 11 nocturnal drizzle days), with no strong tendency
for either category to be clustered in time. The exception
was the grouping of nocturnal drizzle events between
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1211
FIG. 12. As in Fig. 6, but for 0300 UTC (2300 LT 18 Feb) 19 Feb
from Rancho Grande (Ariquemes) sounding location.
FIG. 11. As in Fig. 7, but for 0245 UTC (2245 LT 18 Feb) 19 Feb.
29 January and 4 February, which corresponded to a
weak northwesterly low-level wind regime (Rickenbach
et al. 2002; Halverson et al. 2002). Diurnal composites
and statistics of nocturnal convection and nocturnal
drizzle will be presented in section 3c. However, the
composite analyses obscured important details of the
structure and evolution of these systems that gave insight to their occurrence at night. Thus we turn first to
two illustrative case studies.
1) NOCTURNAL
CONVECTION:
SQUALL LINE SYSTEM
THE 18 FEBRUARY
Afternoon noon-balloon convection was widespread
in the radar area on 17 February and had decayed to
patchy stratiform remnants by 2200 LT (0200 UTC 18
February). Radar and IR imagery indicated that by that
time, a large, preexisting squall line began to enter the
eastern edge of the radar area. Animation of IR imagery
covering Amazonia strongly suggested that this squall
line was a part of a large-scale, westward propagating
line of cloudiness that formed 1.5 days earlier (the afternoon of 16 February) on the northeast coast of Brazil,
of the type described by Cohen et al. (1995). By 1415
UTC 17 February (12 h earlier, late morning local time)
the squall-line system had survived its first night and
was clearly evident on IR imagery (Fig. 4a) moving
westward over central Amazonia, oriented northwest to
southeast and at least 1000 km in length. The line was
identified by a sharp leading edge of high brightness
temperature gradient (the convective line) and cold
cloud trailing with respect to its westward motion (stratiform anvil region), typical of IR signatures of large
squall-line systems (Smull and Houze 1985; Scofield
1987; Zipser 1988; Heymsfield and Fulton 1994; Rickenbach 1999). Twelve hours later, IR imagery at 0215
UTC 18 February (Fig. 4b) showed the squall line still
well defined on its second night, 3000 km west of where
it had formed, with new cold cloud elements on the
southern end entering the TOGA radar view. The leading edge of the cloud line traveled approximately 550
km during this 12-h period (a propagation speed of
roughly 13 m s 21 ), consistent with the speed of coastally
generated lines reported in Cohen et al. (1995). This
event occurred in a regime of low and midlevel easterly
flow in Rondônia (Rickenbach et al. 2002). An easterly
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FIG. 13. As in Fig. 5, but for 19 Feb.
low-level jet has been found to be an important feature
of westward propagating coastal squall lines (Silva Dias
and Ferreira 1992; Warner et al. 2003). Betts et al.
(2002b) implicated this particular nocturnal squall-line
system in raising surface concentrations of ozone at
night in the TRMM-LBA region, by vertical transport
in convective downdrafts. That study also suggested, as
is done herein, that this system originated along the
northeast coast of Brazil.
The daily time series of radar rain area, radar rainfall
rate, radar echo-top height, and IR pixel count (Fig. 5)
captured the evolution of the southern portion of the
squall line beginning at 0000 LT. As the squall line
entered the radar view, stratiform rain area (Fig. 5a) and
IR cloud-top area (250 and 230 K; Fig. 5d) increased.
Radiosonde data from 2300 LT (0300 UTC; Fig. 6a)
indicated an atmosphere primed for organized deep convection, with ample CAPE distributed through the column and an easterly jet centered at 4-km height. The
convection on the leading edge weakened between 0000
and 0300 LT, as shown by the downward trend in the
mean 30-dBZ surface height, convective rain rate, and
convective rain area. This can be seen in the local IR
imagery, radar reflectivity maps (2-km height), and vertical cross sections of radar reflectivity normal to the
squall line (Fig. 7). The vertical cross section at 0015
LT (0415 UTC) showed the northern portion of the line
with deep convection reaching 16-km echo-top height,
and both a trailing and forward-sheared anvil region
(consistent with northeasterly winds in the upper troposphere; Fig. 6a). Two hours later (Fig. 8) at 0215 LT
(0615 UTC) the leading-edge convection had collapsed
in that portion of the line, leaving an extensive stratiform
rain region (11-km-maximum echo top) and large cold
anvil cloud (210 K). During this period the southern
portion of the line remained active. A subsequent resurgence in convective intensity and rainfall between
0300 LT (0700 UTC) and 0600 LT (1000 UTC) seen
in Fig. 5 was traced to the regeneration of the northern
portion of the squall line (Fig. 9). This portion of the
line decayed to extensive and intense stratiform rain
with a strong radar brightband signature (not shown)
MAY 2004
RICKENBACH
FIG. 14. As in Fig. 7, but for 0345 UTC (2345 LT 18 Feb) 19 Feb.
likely due to aggregation of melting snow near the 08C
level, suggestive of the growth of convective debris settling through mesoscale ascent (Rutledge and Houze
1986). By midmorning, all activity had ceased, and the
subsequent noon-balloon convection was suppressed
(Fig. 5), presumably from the stabilizing effect of the
low-level cooling and drying induced by the system
passage (Betts et al. 2002a; postline sounding shown in
Fig. 6b).
2) NOCTURNAL
DRIZZLE:
THE 19 FEBRUARY
EVENT
This event occurred the night following the 18 February squall line discussed in the previous section. By
the evening of 18 February, GOES-8 observed a general
lack of convective activity in the entire western half of
1213
Amazonia, including the TOGA radar area, as shown
in 2145 LT (0145 UTC 19 February) IR image (Fig.
10a). An hour later, IR imagery (Fig. 11a) revealed an
expanding region of midlevel cloudiness (250–240-K
range) over the radar domain. This cloudiness formed
in situ and was not associated with nearby convection.
At the same time the radar indicated the initial formation
of weak echo above the 08C level (Fig. 11b), between
5 and 8 km, within the growing midlevel cloud. In the
center of the cloud mass the radar observations suggested light rainfall at or near the surface (Fig. 11b). A
sounding released from the Rancho Grande site at that
time (Fig. 12) passed through the western edge of this
cloud feature. The sounding profiles indicated two thin
saturated layers in the middle troposphere, one at 580
mb (the 08C level, 4.5 km above the surface) and the
other at about 440 mb (2158C or 7.5 km above the
surface), and a relatively low CAPE value of 858 J kg 21 .
This implied two stratus cloud layers with the lower
cloud base at 4.5 km (the 08C level) and upper cloud
top at 7.5 km (consistent with the IR brightness temperature image in Fig. 11a). Strong southeasterly flow
in the upper troposphere transitioned to weak northeasterly winds at and below cloud top, raising the possibility that shear-induced instability had a role in generating the midlevel cloud. However, such a mechanism
does not explain the nocturnal timing of the cloud formation.
Cloud cover and light precipitation area continued to
increase for the next couple of hours. Time series of
radar and IR parameters (Figs. 5 and 13) in the radar
domain showed that the midlevel cloud cover maximum
(250 K) at 0000 LT (0400 UTC) preceded the rain area
maximum by 1 h. This was consistent with the observed
formation of precipitation (and clouds) in the middle
troposphere and subsequent descent by gravitational settling. Figure 14a shows the cloud cover near the time
of maximum areal extent. Note the expanded area of
low-level stratiform rain and the descent of midlevel
radar echo to the surface (Fig. 14b), particularly in the
northern portion of the domain. Observations from the
National Oceanic and Atmospheric Administration
(NOAA) 915-MHz profiler (location shown in Fig. 1)
captured a similar pattern of evolution that night, with
echo formation between 5 and 7 km AGL, descent and
brightband formation, and light surface rainfall. The
profiler data imagery for this time (which may be seen
online at http://daac.gsfc.nasa.gov/CAMPAIGNpDOCS/
TRMMpFE/lba/profpprelim.shtml) did not indicate evidence of convective upward motion, with 1–2 m s 21
fall speeds above the bright band suggestive of falling
snow. However, the formation of midlevel cloudiness,
height of first radar echo, and presence of the radar
bright band strongly suggested that weak upward air
motion occurred above the melting level.
Two hours later [0145 LT (0545 UTC)], it was clear
from IR imagery (Fig. 15a) that much of the western
Amazon basin was covered by midlevel clouds, which
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MONTHLY WEATHER REVIEW
VOLUME 132
FIG. 15. (a) As in Fig. 4, but for 0545 UTC (0145 LT at the TOGA radar) 19 Feb. (b) As in Fig. 7b, but for
0550 UTC (0150 LT) 19 Feb.
had evolved similarly to the cloudiness over the radar
domain. By then, the cloud deck near the radar began
to gradually dissipate. Only at this later stage, a welldefined radar bright band formed between 3 and 4 km
above the surface, suggesting melting of snow or snow
aggregates. Streamers from the more intense brightband
features coincided with higher values of near-surface
radar reflectivity (Fig. 15b), which suggested that the
melting of snow aggregates enhanced rainfall at the
ground. Surface rainfall data were unavailable for this
event. However, Joss disdrometer data (at the profiler
location; cf. Fig. 1) for the very similar 2 February
nocturnal drizzle event indicated 0.5-mm rain accumulation over 2 h, with a peak rainfall rate of 0.6 mm
MAY 2004
RICKENBACH
1215
occurrence of midlevel cloud at night, though more localized than the four events highlighted in Table 1.
The large-scale extent of this event (Figs. 10 and 15)
argued for a midlevel cloud generation mechanism acting on a similarly large scale. Topographic effects (i.e.,
upslope, orographic lifting) may be discounted by observing that the spatial pattern of cloud formation did
not correspond to the location of increasing terrain slope
in western Amazonia (cf. Fig. 1). Instead, the nocturnal
timing of cloud generation pointed to a radiative mechanism.
It seems likely that the nocturnal squall-line system
the previous morning [discussed in section 3b(1)] played
some role in the formation of this nocturnal drizzle
event, either directly or as a result of the suppression
of the prior afternoon’s convection. In the first case, 12
h had passed between the breakup of the stratiform region of the prior nocturnal system and the onset of nocturnal midlevel cloud formation, with clear conditions
in the interim. Soundings during that period (not shown)
indicated that the midtroposphere dried and warmed.
Thus it was unlikely that subsequent nocturnal midlevel
cloudiness had a direct link to the earlier convective
system. In the second case, the lack of afternoon convection led to clear, moist conditions in the lower troposphere, and increasing relative humidity after sunset.
The first clouds appeared to be shallow and isolated
(western portion of Fig. 11) but rapidly filled in horizontally at and above the 08C level where the precipitation process became active. The speculation was that
the initial low-level clouds formed as radiational cooling
of the surface brought low-level air to saturation. The
reason for their rapid horizontal expansion and precipitation formation just above the 08C level was not clear.
Perhaps the weak mesoscale ascent that likely caused
this expansion was driven by radiative destabilization
of the initial shallow cloud tops, some of which extended to the 08C level.
FIG. 16. As in Fig. 7, but for 0745 UTC (0345 LT) 19 Feb.
h 21 . On 19 February, light rain continued to fall under
a dissipating bright band 2 h later [0345 LT (0745 UTC);
Fig. 16b], associated with the remaining cloud (250–
260-K cloud-top temperatures), as the stratiform cloud
dissipated further (Fig. 16a). The breakup of the nocturnal midlevel cloud was underway over western Amazonia by the approach of sunrise [0545 LT (0945 UTC);
Fig. 10b].
This nocturnal drizzle event was typical of those listed
in Table 1, with the main variation being the timing of
cloud and rain area maxima. The 19 February event
observed by the radar was clearly part of a nocturnal
cloudiness pattern that affected a very large region, essentially the western half of Amazonia. Three other
events clearly showed a similarly broad coverage (see
Table 1), and the other days also suggested widespread
c. Nocturnal precipitation: Composite analysis
Composite diurnal time series for the days with nocturnal convection and nocturnal drizzle are presented in
Figs. 17 and 18. The nocturnal convection composites
(Fig. 17) showed general features similar to those discussed in the 18 February squall-line case study. These
include a rapid buildup to a convective rainfall, echo
top, and 30-dBZ height maximum at 0200 LT, a coincident peak in cloudiness at all IR thresholds, and an
expansion of stratiform rain area to about 0600 LT. The
transition from a majority of convective rain to a majority of stratiform rain occurred by 0500 LT, a time of
decreasing vertical intensity, indicating that deep convection had largely diminished by then. This composite
evolution typified large nocturnal squall-line passage
through the radar area. Such squall-line systems occurred on 7 of the 11 nocturnal convective events (Table
1). Animation of infrared imagery indicated that five of
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MONTHLY WEATHER REVIEW
VOLUME 132
FIG. 17. As in Fig. 5, but composited for nocturnal convection days.
these seven large-scale squall-line systems had their origin 1.5 days earlier on the northeastern coast of Brazil
[section 3b(1)]. Together, these five coastal squall lines
alone accounted for 50% of the total rainfall from the
nocturnal convection category. The remaining two were
traced to large propagating cloud systems that formed
more than a day earlier, one in the Brazilian Highlands
of southeastern Amazonia and the other from extreme
western Amazonia (Fig. 1). The other 4 (of 11) systems
were associated with cloud systems that formed locally
in southwestern Amazonia.
The nocturnal convective systems influenced the timing and intensity of the noon-balloon convection the
following afternoon. The late morning increase in convective rainfall and rain area (Fig. 17) was much less
rapid and was delayed by 2 h (to local noon) compared
to the 49-day composite (Fig. 2). Increased blockage of
solar insolation in the hours after sunrise, as suggested
by twice as much midlevel (250 K) cloud area at the
1000 LT minimum compared to the 49-day composite,
likely contributed to the delay of the noon balloon, in
addition to the low-level cooling and drying by the previous night’s convective overturning [section 3b(1)].
Furthermore, noon-balloon echo-top height and 30-dBZ
surface height for the 49-day mean remained constant
or increased between 1200 and 1800 LT, while for the
afternoon composite following nocturnal convection,
the convective height and intensity decreased after about
1400 LT (Fig. 17c). This suggested that the onset of
deep afternoon convection was delayed, with somewhat
fewer, shallower, and weaker convective cells on the
day following the nocturnal convective systems.
In contrast to the large rainrates associated with nocturnal convection, the main signature of nocturnal drizzle events (Fig. 18) was a peak in the stratiform rain
area near 0200 LT (with a resurgence after 0600 LT).
However, a weak maximum in the 30-dBZ surface
height by 0400 LT, though 2 km lower than that for
nocturnal convection, hinted that weak, embedded convective cells might at times have played a role in the
MAY 2004
RICKENBACH
1217
FIG. 18. As in Fig. 5, but composited for nocturnal drizzle days.
maintenance of nocturnal drizzle, although a strengthening bright band could also explain this observation.
The weak rainfall intensity associated with nocturnal
drizzle events suggested that these systems played a
minor role in the total nocturnal rainfall. Nevertheless,
the expansive midlevel cloudiness associated with these
systems was likely important, for example, to the total
radiation budget. The nocturnal 250-K cloud area maximum was nearly as large (and 2 h earlier) as the nocturnal convection cloudiness. Nocturnal drizzle systems
TABLE 2. Fraction of the 49-day TRMM-LBA campaign rainfall
contributed by nocturnal systems.
Nocturnal convection
Nocturnal drizzle
Total
Total
rain
Covective
rain
Stratiform
rain
25%
3%
28%
20%
1%
21%
31%
4%
35%
contributed only 3% of the 48-day rainfall total (Table
2), while 25% of the total was associated with nocturnal
convective events. The majority of rainfall (72%) was
associated with the afternoon noon-balloon convection.
Moreover, in contrast to the nighttime convection, nocturnal drizzle events apparently had little influence on
the next afternoon’s noon balloon, as seen by the sharp
increase in convective rainrate and rain area in Fig. 18.
4. Conclusions
This paper examined the origins of a secondary nocturnal maximum in cloudiness and precipitation in
southwestern Amazonia during the TRMM-LBA field
campaign, which had been observed previously by many
investigators. The general finding was that following
the collapse of the nearly ubiquitous and locally generated afternoon (‘‘noon balloon’’) convection by 2200
LT, organized deep convection contributed to a post-
1218
MONTHLY WEATHER REVIEW
midnight maximum in raining area and high cloudiness
and to a lesser extent rainfall, following a nocturnal
resurgence in vertical intensity of convection and subsequent expansion of stratiform rain area. For 7 of 11
cases, the timing of the nocturnal convective resurgence
was tied to the propagation of large-scale bands of
cloudiness associated with squall lines generated more
than 2000 km away. Of these seven cases, five could
be traced to large-scale squall-line systems that had
formed 1.5 days earlier along the northeastern coast of
Brazil (e.g., Cohen et al. 1995; Greco et al. 1990). These
five coastal squall lines alone accounted for half of the
rainfall from all nocturnal systems sampled by the radar.
Each of these squall lines occurred in a regime of easterly flow in the lower troposphere (Rickenbach et al.
2002), consistent with the importance of an easterly jet
in the formation and maintenance of the coastal squall
lines (Silva Dias and Ferreira 1992). This scenario was
consistent with the findings of westward propagation in
the phase of diurnal rainfall maxima across the Amazon
basin (Garreaud and Wallace 1997). Nocturnal convective systems were found to delay the onset of and weaken the noon-balloon convection the following afternoon.
Observations suggested that these nocturnal systems delayed the initiation of the next day’s local convection
by cooling and drying from widespread precipitation
and by blocking of solar insolation by extensive morning cloud cover. Moreover, there was a tendency for
decreased 12-h-mean rainfall on the afternoon following
nocturnal convective events. These results suggest the
importance of remote, large-scale, propagating convective systems in modulating the timing and intensity of
the diurnal variation of clouds and rainfall in western
Amazonia.
In addition, a previously undescribed nocturnal drizzle phenomenon, independent from nocturnal convective systems, was shown to contribute significantly to
nocturnal midlevel cloudiness, and slightly to precipitation. On nearly one of every three nights in the 49day sample, widespread cloudiness and precipitation
was generated in situ between 5 and 8 km in height,
suggesting that the origin of this phenomenon lay above
the 08C level and was not rooted in the boundary layer.
Nocturnal midlevel cloudiness was often observed to
form over large portions of western Amazonia in regions
undisturbed during the previous afternoon by deep convective systems. Radar observations showed that the
fallout of ice particles produced a radar brightband signature, suggesting the presence of snow aggregates and
implying mesoscale or large-scale ascent. The generating mechanisms of nocturnal drizzle events were not
known, but observations suggested weak upward vertical motion above the melting level, possibly driven by
nocturnal destabilization of shallow cloud tops. The
greatest significance of nocturnal drizzle events may be
the potential effect on the diurnal radiation budget by
the extensive midlevel nocturnal clouds rather than their
marginal contribution to nocturnal rainfall. Further re-
VOLUME 132
search should be directed at the full life cycles of westward propagating coastal squall lines and on detailed
case studies of nocturnal drizzle events.
Acknowledgments. This project was funded by NASA
Tropical Rainfall Measuring Mission Grant NAG59816-6. Helpful discussions with Drs. Robert Cifelli,
Vanderlei Martins, Steve Nesbitt, Ali Tokay, and Earle
Williams are gratefully acknowledged. The comments
of the reviewers were particularly constructive and improved the paper. Mr. Andrew Negri provided the South
America topography data. Dr. Biswadev Roy produced
the radiosonde data plots.
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