HYDROLOGICAL PROCESSES
Hydrol. Process. 20, 591– 610 (2006)
Published online 26 August 2005 in Wiley InterScience (www.interscience.wiley.com). DOI: 10.1002/hyp.5926
The great Atacama flood of 2001 and its implications for
Andean hydrology
John Houston*
Av. Las Condes, 10373, Of. 60, Santiago, Chile
Abstract:
In February 2001, widespread flooding occurred throughout the Atacama Desert of northern Chile and southern Peru.
It was particularly severe in the Rı́o Loa basin, where roads and bridges were disrupted and the town of Calama
inundated. The instantaneous peak flow in the Rı́o Salado, a tributary of the Rı́o Loa, reached 310 m3 s1 , an order of
magnitude higher than any previously recorded event. The flood is estimated to have a return period of 100–200 years
and is shown to have been caused by intense, long-duration rainfall in the western Cordillera associated with La
Niña. The surface water response is typical of arid areas and highly dependent on antecedent conditions, but is quite
different in perennial and ephemeral catchments. Ephemeral flood flows suffer high transmission losses, recharging
phreatic aquifers. Perennial rivers have lower runoff coefficients, but baseflow levels remained high after the event for
several months due to bank storage rebound and interflow. Extremely high energies of ¾3000 W m2 were generated
by the floods in the Cordillera, becoming less in the Precordillera and downstream. Erosion and sediment transport
were consequently highest in the upper and middle reaches of the rivers, with mixed erosion-deposition in the lowest
reach. The new insights gained from the interpretation and quantification of this event have important implications for
palaeoenvironmental analysis, hazard management, water resource evaluation and the palaeohydrological evolution of
the Andes. Copyright  2005 John Wiley & Sons, Ltd.
KEY WORDS
extreme flood; ENSO; precipitation; runoff; recharge; water resources; Atacama Desert; Andean
geomorphology
INTRODUCTION
The Atacama is perhaps the most intensively hyper-arid desert in the world and has probably been so for much
of the last 10–15 million years (Houston and Hartley, 2003). It is surprising, therefore, to find so many of its
morphological features, both ancient and modern, are formed by the action of fluvial erosion and deposition
(Mortimer, 1980; Naranjo and Paskoff, 1980, 1981).
This contradiction is partly due to the position of the Atacama on the west flank of the Andes, which means
that several perennial and many ephemeral rivers cross this desert, ultimately sourced from precipitation at
high elevation. Deep incision (up to 1200 m) took place over many sections of these rivers during the
Pliocene, 1–3 million years ago (age constraint based on a variety of dated volcanics bracketing incision),
following base-level revision as a consequence of tectonic uplift (Gregory-Wodzicki, 2000) and sea-level
change (Hallam, 1992), as well as fluctuations in climate (Houston and Hartley, 2003).
Subsequently, only relatively minor phases of aggradation and incision have taken place (Rech et al., 2001;
Latorre et al., 2004), with almost all rivers currently existing as underfits in their channels, i.e. bankfull
conditions are met infrequently (Houston, 2002) and rarely exceeded. Such a state, whilst contributing the
bulk of all historical data, is deceiving, and extreme hydrological events may still create major impacts to the
system, despite appearing rarely or never in the historical record.
* Correspondence to: John Houston, Av. Las Condes, 10373, Of. 60, Santiago, Chile. E-mail: [email protected]
Copyright  2005 John Wiley & Sons, Ltd.
Received 5 May 2004
Accepted 9 February 2005
592
J. HOUSTON
This skewed database is exacerbated by the paucity of published analyses of the hydrology of the Atacama
Desert. Some early studies (e.g. Peña, 1970) have been followed by others prepared by the Dirección General
de Aguas (DGA) (DGA, 1987) and aid agencies (UNDP, 1978; JICA, 1995) as well as several on the isotopic
nature of river and groundwater (Aravena and Suzuki, 1990; Aravena, 1995), and there are several unpublished
reports by mining companies on specific aquifers. But, in general, and despite a relatively good monitoring
network for this remote area, there has been little synthesis, and few reliable conceptual frameworks exist. As
a result, several focused studies have been undertaken in the absence of context, leading to less than optimal
interpretations and resource evaluation.
In February 2001, widespread flooding occurred throughout the Atacama Desert of northern Chile and
southern Peru. The flooding was particularly severe in the Rı́o Loa basin where several bridges were ruptured
and the town of Calama inundated. As far as is known, this is the first recorded historical event in the Atacama
Desert of such magnitude and impact.
Here, this extreme event is described from initiation to dissipation within the Rı́o Loa basin, particularly
within the Salado catchment, and an attempt is made to place it in perspective by examining its return period,
cause and impact, as well as the implications for palaeohydrological interpretations, hazard management,
water resources and Andean evolution.
THE RÍO LOA BASIN
Rivers draining the west slope of the Andes between 15 and 30 ° S generally flow directly from east to west,
incised across the largely north–south geological structure (Figure 1) consequent upon the relatively rapid
uplift of the Andes. The only exception to this is the Rı́o Loa, which has subsequently developed significant
north–south reaches, adapted to the local structure. It is not yet clear why the Rı́o Loa is the only river to be
Figure 1. Location map showing the main geomorphic units, rivers and towns (left) and a shaded relief DEM (right) of the Atacama Desert.
The main rivers are (1) Tambo, (2) Osmore, (3) Lluta, (4) Camerones, (5) Loa, (6) Salado, (7) Quebrada Salado, and (8) Copiapo
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
THE GREAT ATACAMA FLOOD OF 2001
593
so adapted, but it does greatly enhance its catchment area, especially in the hydrologically critical area above
3000 m above sea level (a.s.l.).
The Rı́o Loa is a perennial river that crosses the heart of the Atacama Desert and is sourced throughout
most of the year by baseflow from a number of basin aquifers in the high Andes, the Precordillera, the
Calama Basin and the Pampa Tamarugal (part of the Longitudinal Valley). During the wet summer months of
December to March on the Altiplano, multi-peak seasonal floods pass down the river, undergoing transmission
losses that partly replenish the basin aquifers.
The surface water drainage area of the Rı́o Loa is 32 820 km2 , although there is significant area
(>8000 km2 ) draining through the subsurface from the Pampa Tamarugal to the north (Figure 2). Unpublished
hydrogeological data suggest that the subsurface flow from the Pampa Tamarugal to the Rı́o Loa is currently
around 6 ð 106 m3 year1 , but this was probably much greater during the Plio-Pleistocene and would, at
times, have included additional surface water drainage.
The drainage density in the catchment below 3000 m a.s.l. is less than 0Ð1 km km2 , except adjacent to
the main watercourse where densities up to 0Ð2 km km2 are found due to groundwater discharge that creates
side channels which sap upstream, frequently ending in large headcuts up to several hundred metres high with
associated mass slumping. Above 3000 m a.s.l. the drainage densities may reach 0Ð5 km km2 , particularly
on less-permeable Miocene ignimbrites and Palaeozoic to Mesozoic metasediments.
The hypsometric curve (Figure 3) of the basin reveals the dominant geomorphic units that the river crosses.
The peaks, indicative of flatter parts of the catchment, are associated with structurally controlled basins
containing groundwater stored in Neogene and Quaternary sediments and volcanics. Depending on the degree
of connection between the basins and the level of active groundwater storage at any time, such groundwater
may cascade down through the system in the subsurface, as well as provide baseflow to the Rı́o Loa.
The profiles of the Rı́o Loa and its main perennial tributary, the Rı́o Salado (Figure 3), display four major
segments: a very short coastal segment, followed by two relatively well-graded segments relating to the
Longitudinal Valley and Calama basins, and finally complex upper segments associated with the Andean
Cordillera. Since rivers only attain smooth concave profiles where they are capable of removing bedrock at a
rate greater than the uplift rate, the segmented profile is indicative of rapid polycyclic uplift dominating over
erosion (incision).
Figure 2. The Rı́o Loa catchment showing perennial (solid lines) and ephemeral (dashed lines) drainage. The Pampa Tamarugal (part of the
Longitudinal Valley) is now a basin of internal drainage, with some groundwater leakage to the Rı́o Loa basin, but 0Ð5–3 Ma would have
contributed additional surface water
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
594
J. HOUSTON
Figure 3. Morphometry of the Rı́o Loa basin showing a hypsometric curve (left) with the principal basins labelled and the river profile
(right) showing the knick points due to the principal mountain ranges
THE RÍO SALADO SUBCATCHMENTS
The Rı́o Salado is a perennial left-bank tributary of the Rı́o Loa that issues from the Cordillera and flows due
west (Figure 4). At El Sifon, the principal gauging station before its junction with the Rı́o Loa, the catchment
area is 782 km2 , and within this catchment there are a small number of perennial and rather more ephemeral
tributaries.
High-altitude, freshwater springs discharge from the foot of porous Quaternary volcanics (¾0Ð5 m3 s1 ),
which, together with geothermal water discharge from the El Tatio geyser field (¾0Ð6 m3 s1 ), maintain a
mean annual minimum flow of ¾0Ð3 m3 s1 at El Sifon (after abstractions amounting to 0Ð8 m3 s1 ). The
Salado catchment at El Sifon largely drains Miocene to Holocene volcanic rocks situated in the Precordillera
and Cordillera (Figure 4b and Table I).
Figure 4. The Rı́o Salado catchment: (a) the subcatchments where peak flow and precipitation measurements were made (black diamonds
indicate sites reported in Table I, grey sites where supporting data are available but not reported here, and inverted triangles indicate
DGA rain-gauge data reported in Table II); (b) the generalized geology of the catchment (modified from Marinović and Lahsen (1984)
and Houston (2004))
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
595
THE GREAT ATACAMA FLOOD OF 2001
Table I. Morphometric properties of the Salado subcatchments (map nos. refer to Figure 4)
Min.
Max.
Area-weighted Basin Basin
Stream
Map no. Catchment UTM N UTM E Area
Max.
mean
length slope
density
(km)
(km) (km2 ) elevation elevation
stream
(m a.s.l.) (m a.s.l.)
elevation
(km) (m/m) (km km2 ) order
(m a.s.l.)
1
2
3
4
5
6
7
8
9
Salado
Cupo
León
Amarillo
Yalqui
Tuina
Seca north
Seca south
Teca
7535Ð3
7549Ð9
7545Ð5
7541Ð1
7528Ð3
7517Ð1
7507Ð0
7503Ð8
7501Ð2
567Ð3
569Ð8
582Ð8
583Ð3
555Ð1
530Ð0
548Ð0
548Ð2
547Ð0
781Ð6
68Ð5
53Ð3
42Ð7
398Ð5
69Ð4
31Ð5
58Ð4
38Ð5
2950
3150
3350
3300
2700
2750
2800
2880
2850
a Bimodal hypsometric curve, Precordillera dominant, Cordillera
b Unimodal hypsometric curve, Precordillera dominant.
c
5350
3850
4700
4400
3950
3200
3500
3400
3450
3098/3805a
3471b
3997c
3870c
3325b
3306b
3315b
3186b
3194b
43Ð3
13Ð3
12Ð0
13Ð0
32Ð7
21Ð3
18Ð7
15Ð3
18Ð0
0Ð055
0Ð053
0Ð113
0Ð085
0Ð038
0Ð021
0Ð037
0Ð033
0Ð033
0Ð32
0Ð37
0Ð20
0Ð24
0Ð40
0Ð17
0Ð53
0Ð22
0Ð45
4
2
1
1
3
1
1
1
1
present.
Unimodal hypsometric curve, Cordillera dominant.
Two broad types of ephemeral catchment occur: Cupo, León and Amarillo, draining the Cordillera with
relatively steep slopes on Quaternary volcanics, and Yalqui, Tuina, Seca and Teca draining the Precordillera
with less steep slopes on a wider variety of geological units.
Phreatic aquifers occur within the Miocene Jalquinche, Opache and Toconce Formations, respectively
composed of fluvial continental gravels, sands, gypsiferous silts and clays; fluvial calcareous sandstones
and palustrine limestones; and pyroclastic deposits with intercalated fluvial sands and gravels (May, 1997).
The Mesozoic–Palaeozoic basement and the welded ignimbrites of the Miocene Pelon, Sifon and Artola
Formations (de Silva, 1989), are largely impermeable, frequently causing confinement of underlying aquifers.
The phreatic aquifers are generally in contact with the surface water system, as shown by the continuity
of water levels and their similar hydrochemistry (Nazca, 2001a,b). Flow between the surface water and
groundwater systems takes place in both directions, depending on location and stage height. Leakage from
the rivers takes place where they initially cross an aquifer, particularly during flood periods, and drainage
takes place from the lower sections of the aquifers to the rivers, especially during the dry season, thereby
maintaining their perennial nature by baseflow discharge. As a result, the flow profiles of the Rı́os Salado and
Loa are highly variable throughout their length, varying from less than 0Ð3 m3 s1 to more than 1Ð2 m3 s1
during the dry season.
CHARACTERISTICS OF THE REGIONAL CLIMATE
The Atacama Desert owes its hyper-aridity to three principal factors: its zonal location between 15 and 30 ° S
lies in the sub-tropical high-pressure belt, where descending stable air produced by the Hadley circulation
significantly reduces convection, and hence precipitation; the proximity of the Andes upwind inhibits moisture
advection from the east, creating a rain shadow; and the cold Peruvian Current, which upwells along the coast,
inhibits the moisture capacity of onshore winds, creating a persistent inversion that traps any Pacific moisture
below 800 m a.s.l.
The boundary between summer-dominated precipitation towards the north and winter-dominated precipitation towards the south occurs in the central Atacama between 20 and 25 ° S. During the austral summer
(December–March), the Andes experience rainfall when atmospheric circulation, dependent on the location
and intensity of the Bolivian high, allows the advection of moist air from Amazonia (Fuenzalida and Rutllant, 1986; Lenters and Cook, 1997; Garreaud, 1999). Wet episodes tend to occur throughout the western
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
596
J. HOUSTON
Cordillera and Altiplano when strong low-level easterly winds from Amazonia transport adequate moisture
to create saturation during uplift within deep convection cells (Garreaud et al., 2003). Wet episodes lasting
around 5 days tend to be separated by dry periods of around 10 days, with one to three such cycles typically
occurring within the period December to March.
As a consequence of the rain shadow, mean annual precipitation declines rapidly from over 300 mm at
5000 m a.s.l. to 20 mm at 2300 m a.s.l. on the west slope of the Andes (Houston and Hartley, 2003). Below
2300 m a.s.l., associated with the Longitudinal Valley, is a zone of extreme hyper-aridity where mean annual
precipitation is less than 1 mm, due partly to the descending return flow of daytime upslope winds in the
local circulation cell above the inversion layer (Rutllant et al., 2003).
Winter rainfall is largely sourced from northerly and easterly moving frontal systems originating from the
Pacific (Vuille and Ammann, 1997), and within the core area of the Atacama contributes less than 30% of
the mean annual rainfall.
Annual variation in summer precipitation is considerable. The coefficient of variation is generally between
0Ð6 and 1Ð0 above 3000 m a.s.l., but this increases to as much as 2Ð8 in the Longitudinal Valley. Interannual
variation (2–5 years) is considered to be largely due to the influence of the El Niño–Southern Oscillation
(ENSO; Aceituno, 1988; Vuille, 1999; Garreaud and Aceituno, 2001). The warm-phase ENSO (El Niño) tends
to produce dry conditions over the western Cordillera and Altiplano, whereas the reverse is true for the cold
phase ENSO (La Niña). The forcing mechanism appears to act through the collective and interdependent
impact of Pacific sea-surface temperatures (SSTs), the strength of the upper level westerlies and the position
of the Bolivian high.
In addition to ENSO variability, there is a tendency for wet (dry) years to cluster on a decadal scale.
Reid (1987) and Enfield (1992) argue for a link between decadal–centennial variations in SST and solar
activity, Svensmark and Friis-Christensen (1997) suggest reduced solar activity increases global cloud cover,
and Garreaud and Battisti (1999) infer that ENSO and decadal variations are forced by surface heat flux
anomalies, which in turn may be under the influence of solar activity (Labitzke and Matthes, 2003; Oh et al.,
2003). Such ideas, whilst still controversial, might explain a precipitation–solar activity (decadal) correlation
as reported by Clayton (1923) for parts of Chile.
The hyper-aridity of the Atacama Desert is also a result of very high evapotranspiration rates, which
significantly impact the hydrological system. Transpiration losses are effectively limited to areas of riparian
vegetation adjacent to perennial rivers where a broad floodplain exists, whereas evaporation is significant from
flooded areas and soils where the water table is close to the surface. These areas amount to approximately 1%
of the Calama basin and 2Ð5% of the Salado catchment. Evaporation from saturated soils decreases rapidly
with depth to water, usually extinguishing by 1Ð5 m (unpublished micrometeorological and lysimeter data).
Analysis of pan evaporation data from the DGA for 11 stations spread through the northern Chilean Atacama
over the period 1977–95 indicates that interannual variation is considerably less than that of precipitation,
with coefficients of variation ranging from 0Ð05 to 0Ð25, unrelated to elevation. Nevertheless, mean annual
evaporation rates are significantly positively correlated with temperature (r D 0Ð89, p < 0Ð001), and hence
negatively correlated with elevation (r D 0Ð94, p < 0Ð001). The highest rates, over 3500 mm year1 occur
between 1000 and 2000 m a.s.l., declining to less than 2000 mm year1 above 4000 m a.s.l. Clearly, such
rates are only sustainable for short periods of time from open water and are severely reduced when soil
moisture deficits develop during most of the year.
WEATHER CONDITIONS DURING SUMMER 2001
Analysis of daily rainfall at five DGA stations (see Figure 5 and Table II for locations) shows multiple
precipitation events during summer (January–March) 2001 with more than 50 rain-days in the Cordillera,
of which an average of 12 days experienced intensities over 10 mm day1 (Figure 6). The Precordillera,
recorded an average of 26 rain-days (seven over 10 mm day1 ) and 8 days in the Calama basin (two over
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
597
THE GREAT ATACAMA FLOOD OF 2001
Figure 5. Daily precipitation data for selected stations in the Cordillera, Precordillera and Calama basin for summer 2001
Figure 6. Mean annual, summer 2001 and maximum daily intensity (2001) for stations in the Salado catchment at the locations given shown
in Figure 4 and Table II
3 mm day1 ). This is more than triple the mean annual conditions (Figure 6). Of particular note is the fact
that intensities were highest in the Precordillera, whereas durations persisted longer in the Cordillera and thus
helped to fill soil moisture deficits prior to the main event.
The greatest flooding took place on 28 February, and the synoptic events that led to this have been
analysed from data provided by the Dirección Meteorológica de Chile. Upper air charts (200, 500 hPa)
show the Bolivian high to be centred at 13 ° S, 70 ° W prior to 25 February blocking airflows from
Amazonia over the Altiplano. The high moved south on 26 February to 25 ° S and east on 28 February
to 55 ° W, allowing a strong upper level easterly airflow from Amazonia to develop over the Altiplano
and western Cordillera. Thereafter it moved back towards the northwest. During the same period, infrared
images show intense convective activity over the Altiplano and western Cordillera, building from the
north (15 ° S) on 22 February towards the south (25 ° S) by 26 February, reaching maximum intensity on
27 February, covering an area of nearly 1 ð 106 km2 of southeastern Amazonia, the central and southern
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
598
J. HOUSTON
Table II. Summer 2001 precipitation at stations within the Salado catchment (map references refer to Figure 4)
Map
reference
A
B
C
D
E
Station
UTM N
(km)
UTM W
(km)
Elevation
(m a.s.l.)
Mean annual
rainfall
(mm)
Summer 2001
rainfall
(mm)
Max. daily
rainfall
(mm)
El Tatio
Linzor
Toconce
Caspana
Chiu Chiu
7526Ð2
7544Ð7
7539Ð2
7539Ð1
7530Ð2
603Ð0
604Ð8
585Ð9
580Ð7
536Ð0
4320
4096
3350
3260
2524
144
159
96
66
4Ð2
422
406
167
280
20Ð3
21
28
50
35
4Ð5
Figure 7. Surface wind direction and speed at Calama Airport (2260 m a.s.l.) for the period 20 February to 3 March 2001. Wind direction
converted to upslope (easterly, C90° ) or downslope (westerly, 90° ). Dashed lines show maximum daily upslope (upper) and downslope
(lower) velocities. Hatched period relates to storm event of 25–28 February that led to flooding
Altiplano and the western Cordillera of southern Peru and northern Chile, ultimately dying away on 1
March.
Surface wind analysis at Calama (Figure 7) shows the typical pattern of anabatic winds during the
day, i.e. reaching maximum intensities around 1700 h, decaying thereafter with calm conditions around
midnight, followed by a reversal and maximum katabatic winds at around 0800 h. During the 25–28 February
precipitation event, upslope winds were significantly weaker and downslope winds stronger, consistent with the
stronger easterly airflow aloft. Precipitation started on the western Cordillera on the 22 February and generally
built south and west into the Precordillera by 26 February, achieving maximum intensity on 28 February at
elevations of approximately 3500 m a.s.l. Therefore, the data from this event are wholly consistent with
previous studies on the origin of precipitation over the Cordillera (Garreaud et al., 2003) and model studies
conducted by Garreaud (1999) of the circulation mechanisms generating precipitation.
Observations at Chiu Chiu indicate that maximum instantaneous precipitation rates occurred between
2200 and 0300 h on 28 February–1 March, associated with the decay of convective activity, although it
is noteworthy that upslope winds persisted longer than usual throughout that night. This is probably due
to the persistence of some convection cells, as upper level easterly airflows helped them drift west of the
Cordillera crest line and hence produced maximum rainfall intensities in the Precordillera. This mechanism
had significant impact on the time to peak flow and flow volume within the Salado catchment.
Daily catchment precipitation was calculated for January through to March using a point kriging technique
(Cressie, 1991) for the five stations shown in Figure 5 and Table II. The resulting precipitation for the perennial
Salado catchment is shown in Figure 8 and for the ephemeral subcatchments during the 25–28 February event
in Table III. For the Salado catchment there were seven rain-days over 10 mm, culminating in 19Ð8 mm on
28 February and totalling 325 mm for the season. For the other subcatchments, those draining the Cordillera
had nearly twice the event rainfall of those draining the Precordillera (Table III).
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
599
THE GREAT ATACAMA FLOOD OF 2001
Figure 8. Mean daily flow at El Sifon (Rı́o Salado) during summer 2001 compared with catchment rainfall (see text) and separated into its
respective components
Table III. Rainfall and runoff data for the Salado subcatchments for the event of 25–28 February 2001
(map nos. refer to Figure 4)
Map no.
Perennial
1
Ephemeral
2
3
4
Cordillera mean
and standard
deviation:
5
6
7
8
9
Precordillera mean
and standard
deviation:
Catchment
Catchment Rainfall Peak
Unit
rainfall volume flow peak flow
(mm)
(ð106 (m3
(m3
m3 )
s1 ) s1 km2 )
Flood
volume
(ð106 m3 )
Unit
Runoff
flood
(mm
volume
day1 )
(m3 km2 )
Runoff
coefficient
CSW
Salado
49Ð0
38Ð297
310
0Ð396
9Ð305
11 899
11Ð9
0Ð243
Cupo
León
Amarillo
77Ð6
61Ð0
73Ð4
5Ð315
3Ð246
3Ð137
248
236
197
3Ð565
4Ð433
4Ð617
4Ð21 š 0Ð56
0Ð257
0Ð170
0Ð173
3697
3200
4043
3647 š 424
3Ð7
3Ð2
4Ð0
0Ð048
0Ð053
0Ð055
0Ð052 š 0Ð003
Yalqui
Tuina
Seca north
Seca south
Teca
35Ð5
33Ð7
30Ð1
22Ð1
35Ð5
14Ð157
2Ð337
0Ð947
1Ð292
1Ð369
23
39
22
30
25
0Ð057
0Ð559
0Ð709
0Ð522
0Ð659
0Ð50 š 0Ð26
0Ð637
0Ð160
0Ð088
0Ð090
0Ð101
1598
2309
2807
1538
2630
2177 š 584
1Ð6
2Ð3
2Ð8
1Ð5
2Ð6
0Ð045
0Ð069
0Ð093
0Ð069
0Ð074
0Ð070 š 0Ð017
FLOOD FLOWS IN THE RÍO SALADO
Several significant flood flows took place in the Rı́o Salado and Rı́o Loa during January and February 2001,
increasing to nearly 8 m3 s1 on 20 February, but it was not until 28 February that the extreme event took
place that destroyed several bridges (Figure 9) and inundated the town of Calama as a result of overbank
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
600
J. HOUSTON
Figure 9. The Rı́o Loa at Chiu Chiu at 0700 h, 1 March 2001 (3 h after peak flow) showing the ruptured bridge and new channel created
behind the left abutment (a) and 3 months later showing typical flows (b). El Sifon on 1 March showing debris covering the gauging station,
and maximum flood height arrowed (c), flooding in the Pampa Tamarugal, 40 km from the mountain front (d), Quebrada Yalqui showing the
main incised scour channel and braided waning mud flow deposits from the flood (e), and ponding associated with crevasse splay deposits
in the Calama basin, 1 month after the flood (f)
flooding. The event caused erosion, channel reorganization, overbank flooding and deposition throughout the
catchments of the Rı́os Loa and Salado.
An automatic gauging station maintained by the DGA at El Sifon provides mean daily flow data, which
is complete except for the period from 1 March, when it was destroyed by the flood until repaired on 19
Copyright  2005 John Wiley & Sons, Ltd.
Hydrol. Process. 20, 591– 610 (2006)
THE GREAT ATACAMA FLOOD OF 2001
601
March. Missing data were interpolated based on catchment daily rainfall and calibrated against the maximum
instantaneous flow (a) recorded prior to destruction and (b) calculated from surveys of maximum flood levels
and analysed using the slope–area method (Benson, 1968) coupled with the Soil Conservation Service (SCS)
unit hydrograph (US SCS, 1972) which have been shown to give reliable results in northern Chile (Houston,
2002). The resulting hydrograph shows 10 rainfall events associated with an equivalent number of flow peaks
through the season (Figure 8). The calculated maximum instantaneous and mean daily flows occurred on 28
February with rates of 310 m3 s1 and 136 m3 s1 respectively. This compares with typical annual maximum
mean daily flows of 2–8 m3 s1 (1978–2000), with the previous highest recorded being 25 m3 s1 in 1999.
Assuming a simple three-part hydrograph (Horton, 1933; Barnes, 1939), it is possible to separate the
components of baseflow, interflow and overland flow (Figure 8) using an automated technique (Rutledge,
1998). Based on daily data between 1978 and 2000, the master recession curve for the Rı́o Salado at El Sifon
has a baseflow recession constant k of 0Ð997, but for the period October–December 2000 it was 0Ð995, due
to the prior wet year (1999) and consequent higher initial groundwater (baseflow discharge) levels, indicating
non-linear storage reservoirs.
The lag time between the centroid of a rainfall event and the associated flood peak at El Sifon is plotted
in Figure 10a. This indicates that for the seven events leading up to the flood of 28 February the lag time
reduced from 7 days to 1 day, indicating the importance of antecedent conditions and the progressive refilling
of soil moisture storage. Thereafter, lag time remains stable due to storage having been fully recharged. The
observed lag time for the Rı́o Salado for the event of 28 February was approximately 5 h, compared with
6Ð2 h determined from the SCS method, reaffirming the appropriateness of its use in this environment.
Figure 10. Lag time of the flood event through the catchment: (a) rainfall centroid to peak flow, (b) peak flow at El Sifon to peak flow at
mouth of Rı́o Loa (325 km), (c) the residual flood in the river system after ¾40 days (April) and ¾80 days (May) and (d) flood velocity
versus magnitude between El Sifon and the sea. In (b) the last recorded event of 28 February destroyed the gauging station at the Rı́o Loa
mouth, and hence there are no later data
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Hydrol. Process. 20, 591– 610 (2006)
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J. HOUSTON
The passage of successive overland-flow flood waves through the catchment from El Sifon to the sea at
Desembocadura (sea mouth) determined from DGA flow records, a distance of 325 km, depends either on
antecedent conditions (Figure 10b) or flood magnitude (Figure 10d), or possibly both. Ben-Zvi et al. (1991)
showed that flood wave velocities were related to magnitude in Israel. In the Rı́o Loa, the first flood peak at
El Sifon did not transfer to the sea, since it was relatively insignificant and was entirely dissipated. Thereafter,
a progressive decrease in lag time from 11 days to 3 days occurred, giving increasing flood wave velocities
of 0Ð34 to 1Ð25 m s1 as flood magnitudes increased. However, the magnitudes of the flood peaks at the
sea were all similarly small (typically 0Ð04 m3 s1 , or 30%, relative to background) until the flood event
of 28 February, which was 0Ð31 m3 s1 (0Ð23 m3 s1 above background or 290%), and an equally valid
interpretation would be that the increase in flood wave velocity is due to reducing transmission losses as
storage reservoirs (aquifers) become recharged.
Subsequent spot gauging throughout the catchment during April and May (Figure 10c) shows the passage
of a residual flood wave through the catchment, interpreted to be bank storage rebound and interflow. Based
on the two peaks in April and May, the velocity of this wave is 0Ð02 m s1 and appears to be rapidly damped
downstream from the area of highest infiltration in the western Cordillera, Turi and Yalqui basins (Figure 3).
EVENT FLOWS IN EPHEMERAL SUBCATCHMENTS
Only minor flow took place in the ephemeral subcatchments of the Rı́o Salado prior to the event of 25–28
February. This event, however, created significant flow in virtually all subcatchments draining the Cordillera
and Precordillera, even extending right across the Longitudinal Valley farther to the north (Figure 9d). Peak
flows were recorded by flood debris in the upper parts of these catchments, and overbank splays produced
large areas of ponding on the lower playa areas, which lasted up to 8 weeks.
Flood debris allowed the identification of peak flow levels, which were subsequently surveyed and maximum
instantaneous flows calculated using the slope–area method, and SCS unit hydrographs allowed flood volumes
to be determined. These values are given in Table III. Absolute and unit peak flows are primarily dependent
on area, which explains 76% of the variance (p < 0Ð01). Peak flows are much higher for subcatchments in the
Cordillera than in the Precordillera, as would be expected (elevation explains 34% of the variance, p < 0Ð1),
and unit peak flow did not differ significantly between perennial (the Salado is dominated by drainage from
the Precordillera) and ephemeral catchments draining the Precordillera.
On the other hand, unit flood volume for the perennial catchments was four times greater than for the
ephemeral catchments, regardless of elevation. Thus, runoff coefficients for perennial and ephemeral rivers are
completely distinct (Table III, Figure 11). Runoff coefficients for each type of catchment show a remarkable
linear relationship between rainfall and runoff (explaining 99% of the variance, p < 0Ð001 for both perennial
and ephemeral catchments), despite draining a wide variety of geomorphological and geological environments
(Figure 4). No significant relationships were found with any geomorphic variables, such as basin slope, stream
density or order, calling into question whether the geomorphic unit hydrograph concept (Rodrı́guez-Iturbe and
Valdes, 1979; Gupta et al., 1980) would be applicable in this environment. It is perhaps surprising that the
runoff relationships from the group of catchments draining Quaternary volcanics in the Cordillera (Cupo, Leon
and Amarillo) are so similar to those draining Mesozoic–Palaeozoic bedrock and Miocene volcanics of the
Precordillera (Yalqui, Tuina, Seca and Teca). However, it is possible that this is a result of the measured flow
sections all being located close to the mountain front–pediment junction, where infiltration is concentrated,
and thus runoff at these sites being controlled more by local than catchment infiltration characteristics.
Runoff coefficients CSW are also inversely related to catchment area (see Tables I and II and Figure 11),
for perennial rivers increasing from 24% (782 km2 ) to 40% (15 km2 ) and for ephemeral catchments from
4Ð5% (399 km2 ) to 9Ð3% (32 km2 ).
The very different flood magnitudes and runoff coefficients for perennial and ephemeral rivers is strongly
indicative of extremely high transmission losses (aquifer recharge) in the latter catchments. In contrast, the
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Hydrol. Process. 20, 591– 610 (2006)
603
THE GREAT ATACAMA FLOOD OF 2001
Figure 11. Runoff volume plotted against rainfall volume for stations in Table III (solid) with additional data not provided herein (open).
Figures attached to data points are runoff coefficients
higher magnitudes in perennial rivers may be attributed to their greater incision, which causes them to be in
constant contact with the groundwater phreatic surface, hence allowing perennial discharge and limiting flood
flow infiltration capacity in materials that are already partially saturated.
EVALUATION OF THE RETURN PERIOD
Extreme events are notoriously difficult to analyse, especially if records are short and climatic non-stationarity
is suspected (Kochel and Baker, 1982; US NRC, 1999). The log Pearson type III (LP3) distribution can,
however, provide acceptable fits to hydrological extremes in a wide variety of situations (US WRC, 1981;
López et al., 2002) and was used to estimate return periods for rainfall and Rı́o Salado flows at El Sifon
(Table IV); no data exist to estimate return flows for ephemeral streams.
Table IV. Return periods associated with rainfall and runoff for summer 2001 based on 23 years data (1977–2000)
Rainfall intensity
El Tatio
Linzor
Toconce
Caspana
Chiu Chiu
El Sifon
Mean š SD
Rain days
>10 mm day1
Daily
Monthly
Annual
8
11
70
45
n/a
12
14
15
15
14
20
20
14
35
20
65
60
8
25
n/a
34 š 30
14 š 1
22 š 8
40 š 28
Copyright  2005 John Wiley & Sons, Ltd.
Runoff
Daily
Monthly
Annual
120
150
170
Hydrol. Process. 20, 591– 610 (2006)
604
J. HOUSTON
Daily maximum rainfall return periods vary considerably, suggesting that, although the storm event of
25–28 February was widespread throughout the Altiplano and northern Chile, multi-cell convection intensity
was locally variable, partly as a result of invigoration by upslope wind convergence within the topographic
amphitheatre of the Turi Basin and partly as a result of internal cell dynamics creating local downbursts with
intense precipitation (typically associated with convective decay, as noted above).
Annual return periods are greater than monthly return periods due to the longevity of precipitation
throughout the summer of 2001, and this longevity is unusual, as shown by the high return periods for
the number of rain-days over 10 mm, especially in the Cordillera.
The extreme flood flows at El Sifon on the Rı́o Salado during 2001 tend to give equivocal fits to an LP3
distribution. The best estimates are given in Table IV, and suggest that the flows had a return period of
between 100 and 200 years. Under any circumstances, the non-linearity of the rainfall-runoff process is thus
evident.
AQUIFER RECHARGE DUE TO FLOOD WATER INFILTRATION
Phreatic aquifers occur in the Opache and Toconce Formations of the Calama and Turi basins respectively.
These aquifers have typical saturated thicknesses from 50 to 100 m, and are in hydraulic contact with the
perennial Rı́os Salado and Loa. Monitoring wells penetrating these aquifers record the flood event of summer
2001 as recharge, initiated immediately after the flood event and peaking around May 2001, some 3 months
later, as a result of the buffering effect of the groundwater reservoir (Figure 12). The rise in hydrograph is
widespread through the basins, but is of variable magnitude due to aquifer and recharge heterogeneity. Since
the hydraulic gradient of the phreatic surface is relatively high (0Ð008–0Ð012), it can be assumed that recharge
over large areas distant from the perennial rivers is due to either direct precipitation infiltration or flood flow
infiltration from ephemeral streams.
The volume of recharge can be estimated by making a number of assumptions:
1. The majority of recharge occurs close to the mountain front, where ephemeral rivers discharge onto the
pediment surface.
2. Recharge at the mountain front is transmitted through the aquifer.
3. The recharge coefficient CGW D 1 CSW .
4. The measured subcatchments are representative of non-measured areas.
Assumptions 1 and 2 were demonstrated by Houston (2002) to be true for the Pampa Tamarugal in the
Longitudinal Valley of northern Chile and are generally accepted as common in arid zones (Simmers, 1997).
Assumption 3 is considered reasonable because: (a) all flow downstream of the measurement points (close
to the mountain front) can be conservatively assumed to evaporate, so that (1 CSW ) represents all water
precipitated over the catchment that did not flow past the measuring point; (b) evaporation losses above the
measuring point are effectively zero, because nearly 90% of the flood in the ephemeral channels occurred in a
single night, whilst cloudy convective conditions prevailed, preventing insolation and lowering temperatures
(Garreaud, 2000), and, as a result, (1 CSW ) becomes equivalent to groundwater recharge. Assumption 4 is
also reasonable, given the similarity of catchment response demonstrated previously.
As a result, the total volume of recharge for the two catchments distant from the perennial rivers can
be estimated (Table V). This can be checked by making predictions of the phreatic water level (GWL) rise
based on assumed values of specific yield (Sy) for the aquifers and comparing with observed GWL rise
in the monitoring wells. In both basins it can be seen that the mean observed GWL rise is close to that
predicted for an Sy of 1Ð0–1Ð5 ð 101 . Pumping tests on wells in the Turi basin give a range of values for
Sy of 0Ð3–1Ð2 ð 101 and in the Calama basin of 0Ð7–1Ð5 ð 101 , thereby confirming the validity of the
estimates.
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Hydrol. Process. 20, 591– 610 (2006)
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THE GREAT ATACAMA FLOOD OF 2001
Figure 12. Groundwater response to the flood event. Hydrographs from wells in the Calama and Turi basins (above) show the impact of
recharge. Map of well locations and generalized aquifer zones for which recharge was calculated (below) giving the observed GWL rise in
October 2001 as extrapolated from the prior recession. Also shown are contours (100 m interval) on the surface of the phreatic water table
Table V. Recharge estimates due to event of 25–28 February 2001
Basin
Turi
Calamaa
a
Groundwater recharge
volume (ð106 m3 )
Event recharge
coefficient CGW
33Ð3
31Ð6
0Ð949
0Ð946
Predicted GWL rise based on
assumed Sy values (m)
Sy D 0Ð05
Sy D 0Ð10
Sy D 0Ð15
1Ð48
0Ð97
0Ð74
0Ð49
0Ð49
0Ð33
Observed mean
GWL rise (m)
0Ð65
0Ð34
Southeast sector only; see grey area on Figure 12.
CAUSES OF THE FLOOD
The origin of the 2001 flood lies primarily in the weather conditions during summer 2001: (a) extensive rainfall
over the western Cordillera over an exceptionally long period and (b) intense rainfall in the Precordillera, and
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Hydrol. Process. 20, 591– 610 (2006)
606
J. HOUSTON
Figure 13. The 2001 event in context: long-term rainfall, surface flows and groundwater levels compared with solar and ENSO activity
(from NOAA-CPC, National Geophysical Data Center archive)
especially in the Rı́o Salado catchment. These conditions occurred at the end of 3 years of predominantly
cold-phase ENSO (La Niña; Figure 13). The only other occurrence of a 3-year La Niña since 1970 was
1973–75, which also led to intense rainfall in the western Cordillera, but there are no hydrological data
available and no anecdotal records of severe flooding.
Since 1970, all above-average precipitation in the western Cordillera occurred whilst La Niña conditions
prevailed in the same summer or previous winter. However, only 75% of cold-phase ENSO years lead to
high rainfall, with the remaining years apparently showing less intense or shorter duration SST anomalies in
the eastern Pacific (area El Niño 1 C 2).
There is no strong evidence for any linkage between decadal variations in solar activity and ENSO, as
suggested by Reid (1987), or with precipitation, as suggested by Clayton (1923).
Long-term flows in the Rı́o Salado at El Sifon (Figure 13) are correlated with mean annual rainfall
(r D 0Ð571, p < 0Ð01), as would be expected, but only 60% of wet years give rise to peak flows significantly
greater than background, suggesting that antecedent soil moisture conditions (rainfall duration) play an
important role. Conditions during 2001 certainly point to this factor as being of crucial importance.
Monthly flow conditions in the Rı́o Salado are complex, however, with seasonal variations in baseflow
being marked.
Groundwater levels are dominated by seasonal variations in the order 0Ð1–0Ð4 m year1 , compared with
long term variations of 0Ð01–0Ð03 m year1 . Long-term variations in wells T2, C1 and CC3 do, however, show
a tendency to rise during the period 1991–2001 in parallel with annual rainfall during the same period. Well
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Hydrol. Process. 20, 591– 610 (2006)
607
THE GREAT ATACAMA FLOOD OF 2001
T7 suggests that GWLs were still in long-term recession during 1990–96, before responding to the increased
rainfall by the mid–late 1990s. This may indicate that the source of recharge to the aquifer penetrated by T7
is distant. All wells responded relatively dramatically to the 2001 event.
Therefore, both surface and groundwater respond to long-term and short-term variations in precipitation;
but the relationship is not simple, and the impact of additional controlling factors, such as aquifer compaction,
seasonal variations in evaporation and winter snowfall/snowmelt, requires further investigation.
GEOMORPHOLOGICAL IMPACTS OF THE FLOOD
The flood of 28 February generated high energies in the ephemeral channels; a maximum of >5000 W m2
(mean ¾3000 W m2 ) in the Cordillera, and between 100 and 1000 W m2 in the Precordillera. At El Sifon,
on the Rı́o Salado, the maximum specific stream power was 425 W m2 . Such high energies necessarily
caused considerable erosion and channel change (Figure 9). In the Turi basin, several new distributary channels
formed on the pediment-playa surfaces with incisions up to 2 m depth and 5 m width, causing widespread
disruption to gravel roads.
The perennial Rı́os Loa and Salado have high-energy, confined floodplains (Nanson and Croke, 1992)
throughout considerable sections of their length; but, even where confined by ‘bedrock’, channels changed
location (Figure 9a), causing considerable damage. In the perennial rivers, extreme levels of suspended
sediment and bedload exacerbated scour and incision; and although erosion dominated the upper and middle
reaches, even in the lower reaches bank erosion took place, alternating with aggradation of sand and gravel
bars and crevasse splays, with considerable volumes of sediment being evacuated from the catchment. At
Desembocadura, specific stream power was estimated to have declined to less than 5 W m2 .
Overbank flooding and mudflows occurred almost universally in both ephemeral and perennial catchments,
creating large areas of standing water within basin playas and in the town of Calama, and ultimately depositing
large areas (hundreds of square kilometres) and thicknesses (to ¾1 m) of fine-grained sediments. In sections
where the channel was not well confined or was braided, incision in part of the channel was balanced by
overbank deposition and the consequent generation of terraces (Figure 9e), which would appear to be typical
for floods in such environments (Hooke and Mant, 2002).
Based on the known extent of the flood deposits and 15 profiles, Quebrada Tuina is estimated to have
redistributed up to 5 km3 of sediment from the Precordillera to the Calama basin playa (Figure 9f), representing
¾70 mm erosion over the catchment (assuming similar densities for eroded and deposited materials, which is
not unreasonable considering the Precordillera catchments have a widespread cover of cohesionless aridisols).
IMPLICATIONS FOR PALAEOENVIRONMENTAL INTERPRETATION, HAZARD MANAGEMENT
AND WATER RESOURCES
Flooding in the Atacama Desert is generally associated with El Niño (Glantz, 1996; Vargas et al., 2000;
Keefer et al., 2003). However, as Dettinger et al. (2000) suggest, the reality is more complex. Flooding
associated with El Niño is largely a result of endogenous winter precipitation and is limited to the Coastal
Cordillera, where the impact of extra-tropical depressions is felt. Flooding in the Atacama Desert (including the
Longitudinal Valley, Precordillera and Cordillera), on the contrary, is largely due to exogenous precipitation
associated with La Niña, as demonstrated by this study and Houston (2002). Consequently, the interpretation
of palaeohydrological events and Quaternary geology need careful assessment of the source of floodwater
before correlations with either El Niño or La Niña are inferred.
Similarly, climatically or hydrologically induced hazards in the Atacama Desert require evaluation based
on different circumstances, depending on whether they are El Niño (winter) or La Niña (summer) induced.
Extreme El Niño conditions do not lead to debris flows or mud slides in the interior of the Atacama Desert,
only in coastal towns such as Antofagasta, whereas extreme La Niña conditions may lead to flooding, erosion
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J. HOUSTON
and communication difficulties in the interior and coastal towns such as Arica, where exogenous perennial
rivers reach the coast. This has important implications for strategic emergency planning, as well as for the
management of hydraulic structures, such as control and impoundment reservoirs.
With regard to water resources, it is evident that ‘average’ conditions do not exist, and the evaluation and
interpretation of many hydrological studies is flawed as a result. Extreme events, such as the one described,
exert a disproportionate impact on the system. Groundwater model simulations, now considered the sine qua
non of water resource evaluations in northern Chile, universally use average conditions as input; but, given
the extreme variability of recharge and the non-linearity of the storage reservoirs, this renders their output
questionable and their predictions subject to error. Unfortunately, there is rarely an adequate database to
evaluate resources properly, especially given the non-stationarity of forcing factors. Nevertheless, a deeper
understanding of the mechanisms leading to process variability can help to provide more realistic constraints
on water resource evaluation.
CONCLUSIONS
In contrast to many arid areas, precipitation over the Altiplano and western Cordillera tends to be spatially
coherent and has been shown to be at least partially forced by ENSO activity. Nevertheless, local factors, such
as boundary-layer flow convergence caused by the Andean topography and convection cell location/timing,
create mesoscale variability that is intensified with decreasing elevation and is amplified by varying catchment
characteristics.
At elevations above 2500 m a.s.l., particularly above 4000 m a.s.l., precipitation enters the hydrological
cycle directly as groundwater recharge (Nazca, 2001a). This cascades down the western slope of the Andes
through a system of interconnected basin aquifers, as well as generating overland flow during high-intensitymagnitude events.
Return periods for significant flooding and groundwater recharge events tend to group into centennial (the
2001 event) and decadal (Houston, 2002) periods. The high energies associated with such floods demonstrate
their catastrophic nature and imply metastable equilibrium (Graf, 1988) and correspondingly long relaxation
times (Knighton, 1998).
In many respects, Atacama rivers respond in a similar fashion to those in other arid regions: flood magnitudes
and peak flows are proportional to catchment area and unit peak flow is dependent on catchment elevation
via rainfall and slope. Furthermore, the magnitude and response time of flooding are strongly controlled by
antecedent conditions.
However, in the Atacama Desert there appears to be surprisingly little geological control on runoff processes.
Much more important is the distinction between perennial and ephemeral rivers. Ephemeral rivers tend to be
much less deeply incised, and discharge to pediments and playas associated with basins of complete or partial
internal drainage. They have low runoff coefficients and generate significant groundwater recharge as a result
of high transmission losses. This groundwater is subsequently transmitted through the volcano-sedimentary
basin aquifers, moving under gravity in the subsurface and partly emerging as baseflow in perennial rivers.
Perennial rivers are deeply incised, crossing geological structures and discharging to the sea, and have higher
runoff coefficients.
Under current conditions, which have probably persisted since at least the Miocene (Houston and Hartley,
2003), water transport and erosion seems to be a two-stage process. In basins dominated by internal drainage
there is flow/erosion from mountain zones and groundwater recharge/deposition in the basins. Cutting across
such basins and the geological structure are deeply incised canyons that have been created partly as a result of
tectonically induced headwater erosion, coupled with exogenously generated surface water floods, and these
act as drainage conduits from the Andean basins.
Thus, the analysis of the 2001 flood provides support for the theory of Hoke et al. (2002), that the large
quebradas (incised rivers) draining the western Cordillera are groundwater generated, and leads to the concept
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Hydrol. Process. 20, 591– 610 (2006)
THE GREAT ATACAMA FLOOD OF 2001
609
that the intermontane sedimentary basins of the Andes are slowly being dewatered over geological time as a
result of tectonic uplift, base-level revision and mainstream incision by groundwater sapping.
ACKNOWLEDGEMENTS
Nazca S.A. provided the funding for this study. The Dirección Meteorológica and Dirección General de Aguas
de Chile provided data.
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