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The origin of brines and salts in Chilean salars: a hydrochemical review
Earth-Science Reviews 63 (2003) 249 – 293 www.elsevier.com/locate/earscirev
The origin of brines and salts in Chilean salars: a hydrochemical review Francßois Risacher a,*, Hugo Alonso b, Carlos Salazar c b
a CNRS-IRD, Centre de Ge´ochimie de la Surface, 1 rue Blessig, 67000 Strasbourg, France Universidad Cato´lica del Norte, Departamento de Quı´mica, Angamos 0610, Antofagasta, Chile c Direccio´n General de Aguas, Ministerio de Obras Pu´blicas, Morande´ 248, Santiago, Chile
Received 26 June 2002; accepted 20 February 2003
Abstract Northern Chile is characterized by a succession of north – south-trending ranges and basins occupied by numerous saline lakes and salt crusts, collectively called salars. Fossil salt crusts are found to the west in the extremely arid Central Valley, while active salars receiving permanent inflows fill many intravolcanic basins to the east in the semiarid Cordillera. Sea salts and desert dust are blown eastward over the Cordillera, where they constitute an appreciable fraction of the solute load of very dilute waters (salt content < 0.1 g/l). The weathering of volcanic rocks contributes most components to inflow waters with salt content ranging from 0.1 to 0.6 g/l. However, the average salt content of all inflows is much higher: about 3.2 g/l. Chemical composition, Cl/Br ratio, and 18O – 2H isotope contents point to the mixing of very dilute meteoric waters with present lake brines for the origin of saline inflows. Ancient gypsum in deep sedimentary formations seems to be the only evaporitic mineral recycled in present salars. Saline lakes and subsurface brines are under steady-state regime. The average residence time of conservative components ranges from a few years to some thousands years, which indicates a permanent leakage of the brines through bottom sediments. The infiltrating brines are recycled in the hydrologic system where they mix with dilute meteoric waters. High heat flow is the likely driving force that moves the deep waters in this magmatic arc region. Active Chilean salars cannot be considered as terminal lakes nor, strictly speaking, as closed basin lakes. Almost all incoming salts leave the basin and are transported elsewhere. Moreover, the dissolution of fossil salt crusts in some active salars also carries away important fluxes of components in percolating brines. Evaporative concentration of inflow waters leads to sulfate-rich or calcium-rich, near-neutral brines. Alkaline brines are almost completely lacking. The alkalinity/calcium ratio of inflow waters is lowered by the oxidation of native sulfur (reducing alkalinity) and the deposition of eolian gypsum (increasing Ca concentration). Theoretically, SO4-rich inflow waters and their derived SO4-rich brines should be found in the intravolcanic basins of the Cordillera because of the ubiquity of native sulfur, while Ca-rich brines should prevail in sedimentary basins where Ca-rich minerals are abundant. This relation is perfectly observed in the salar de Atacama, the largest in Chile. However, several salars located within the volcanic Cordillera belong to the Ca-rich group. Inflows and brines may have acquired their Ca-rich composition in Pleistocene time when their drainage basins were mainly sedimentary. Later on, recent lava flows and ignimbrites covered the sedimentary formations. Underground waters may have kept their early sedimentary signature by continuous
* Corresponding author. Tel.: +33-3-90240409; fax: +33-3-88367235. E-mail addresses: [email protected] (F. Risacher), [email protected] (H. Alonso), [email protected] (C. Salazar). 0012-8252/$ - see front matter D 2003 Elsevier Science B.V. All rights reserved. doi:10.1016/S0012-8252(03)00037-0
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F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
recycling. However, the weathering of volcanic rocks tend to slowly shift the water compositions from the Ca-rich to the SO4-rich type. D 2003 Elsevier Science B.V. All rights reserved. Keywords: Northern Chile; Hydrochemistry; Closed basin; Salar; Salt recycling; Brine evolution
1. Introduction The Central Andes in Chile, Bolivia and Argentina contains a large number of closed basins whose central depression is occupied by saline lakes and salt crusts (salars). Active research has mainly been focused on the world’s largest salt crust: the salar de Uyuni in the Bolivian Altiplano (Ericksen et al., 1978; Rettig et al., 1980; Risacher and Fritz, 1991a, 2000; Fornari et al., 2001; Baker et al., 2001). Information on other small salars are not easily available. They can be found with difficulty in many technical reports, theses and communications at local symposia and bear mostly on economic geology and water resources of the basins. Nevertheless, Stoertz and Ericksen (1974) and Chong (1984) described in detail the features of the salar crusts of northern Chile, but gave little hydrochemical information. Risacher and Fritz (1991b) studied the geochemistry of most of the small intravolcanic salars of the southern Bolivian Altiplano. The origin of solutes in a few Chilean salars has been investigated by Alpers and Whittemore (1990), Spiro and Chong (1996), Alonso and Risacher (1996), Garce´s (2000a,b) and Carmona et al. (2000). Other studies of Chilean salars deal with paleohydrography and paleoclimatology (Grosjean, 1994; Valero-Garce´s et al., 1996; Bao et al., 1999; Grosjean et al., 2001), sedimentology (Valero-Garce´s et al., 1999) and water quality (Caceres et al., 1992; Aguirre and Clavejo, 2000). In this paper, we present a review of the hydrogeochemistry of almost all active salars of northern Chile: 52 saline lakes and salt crusts are distributed along a 1000 km north – south transect in the Andes Cordillera. The drainage basin of an important river and its tributaries (Rio Lauca), feeding salar de Coipasa in the Bolivian Altiplano, has been included. We do not undertake a detailed geological, mineralogical, sedimentological and paleolimnological study of each salar, but to review as a whole the major
hydrogeochemical trends, focusing on the origin of salts and brine evolution. This approach is possible because all these salars belong to the same geographical, geological and climatological environment in spite of their scattering over a 200,000 km2 area. The large number of basins and water samples smoothes local irregularities and provides a global view of their chemical trends. Fig. 1 presents the study area and the location of all salars. Each basin or salar is identified in the text by a number (from Fig. 1) bracketed after its name or abbreviation. Table 1 gives the detailed location and the physiographic and climatic features of the basins. All minerals cited in the text are listed with their formula in Table 2.
2. Analytical data This study is largely based on the analyses of 354 inflow waters and 227 brackish lake waters and brines of Chilean salars (Risacher et al., 1999). However, bromine is not included in this report. Inasmuch as the Cl/Br ratio is an important clue to the origin of solutes, we present in Appendix A the previously unreported Br analyses together with Cl concentration. Both components were analysed by ion chromatography (Dionex 4000i). A discussion of the hydrogeochemical setting of each salar can be found in Risacher et al. (1999). For the giant salar de Atacama, we rely on the comprehensive studies of Moraga et al. (1974, 206 analyses) and Ide (1978, 389 analyses).
3. Regional geology Northern Chile is characterized by contractional structures and intensive block faulting due to the subduction of the Nazca Plate below the South American Plate. A succession of north – south-trending
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
251
Fig. 1. Map of northern Chile and topographic sections showing the main morphostructural units and the location of salars. The elongated and thin formation in the Pre-Andean Depression, between salar de Atacama and Domeyko Range, is the Cordillera de la Sal.
252
Table 1 Physiographic and climatic features of closed basins in northern Chile Code
LAGUNAS COTACOTANI LAGO CHUNGARA RIO LAUCA SALAR DE SURIRE SALAR DE PINTADOS LAGUNA LAGUNILLA SALAR DEL HUASCO SALAR DE COPOSA SALAR DE MICHINCHA SALAR DE ALCONCHA SALAR DE CARCOTE SALAR DE ´N ASCOTA SALAR DE ATACAMA SALAR DE TARA LAGUNA HELADA SALAR DE AGUAS CALIENTES 1 SALAR DE PUJSA SALAR DE LOYOQUES LAGUNA TRINCHERA
COT
No. (Fig. 1)
Latitude (South)
Longitude (West)
Altitude Salar (m)
Alt max (m)
1
18j11V08
69j13V11
4495
6342
CHR
2
18j14V48
69j09V13
4530
LAA SUR
3 4
18j27V00 18j50V03
69j14V54 69j02V08
PIN
5
20j31V38
LGU
6
HCO
Temperature mean (jC)
Prec (mm/ year)
Evap (mm/ year)
TDS min (mg/l)
TDS max (mg/l)
1.9
379
1070
299
2002
SO4
119
6
6
6342
1.9
338
1230
47
1633
SO4
273
22.5
22.5
3892 4260
6063 5780
4.2 2.7
370 250
1200 1280
108 108
784 285,000
SO4 SO4
2374 574
– 144
– 9.5
69j40V17
980
5785
18.5
75
1630
199
365,621
SO4
17,150
377
0
19j55V56
68j50V40
3900
5190
4.6
150
1490
267
SO4
129
7
20j18V18
68j50V22
3778
5190
5
150
1260
108
113,093
SO4
1572
51
2.5
COP
8
20j38V12
68j39V33
3730
5220
5
150
1300
119
330,671
SO4
1116
85
5
MIC
9
20j58V48
68j33V09
4125
5407
3.5
200
1620
94
62,662
SO4
282
2.5
0.6
ALC
10
21j03V07
68j29V14
4250
5407
3.5
200
1620
116
99,361
SO4
128
3.8
0.75
CAR
11
21j22V36
68j20V55
3690
6176
5.8
125
1630
88
335,536
SO4 – Ca
561
108
ASC
12
21j33V01
68j16V58
3716
6023
5.8
125
1630
89
119,853
SO4 – Ca
1757
243
ATA
13
23j25V05
68j15V57
2300
6233
14
160
1800
243
339,719
SO4 – Ca
18,100
3000
TAR
14
23j01V59
67j16V30
4400
5816
0
150
1500
287
176,486
SO4
2035
48
HEL
15
23j05V58
67j08V15
4300
5745
0
180
1500
118
308,546
SO4
221
AC1
16
23j07V22
67j24V38
4280
5370
1
150
1500
1359
122,890
Ca
281
15
2.5
PSA
17
23j12V26
67j30V35
4500
6046
1
150
1500
239
57,578
SO4
634
18
5
LOY
18
23j14V36
67j16V55
4150
5370
1
150
1500
163
242,456
Ca
676
80
5
TRI
19
23j24V01
67j23V51
4290
5060
0
200
1500
10,830
10,830
SO4
59
1276.0
Evol. path
Area basin (km2)
Area Salar (km2)
0.2
5.8
0.4
Area water (km2)
0.15
3.5 18 12.6 14 5.8
0.3
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
Basin
MUE
20
23j26V08
67j26V28
4295
5109
0
200
1500
30,916
97,153
SO4
41
AC2
21
23j28V51
67j33V11
4200
6046
1
150
1500
2530
13,656
Ca
1168
LEJ MIS
22 23
23j29V44 23j43V19
67j41V28 67j45V58
4325 4120
5924 5910
1 2
150 180
1500 1500
1273 217
70,723 5308
SO4 SO4
193 303
MIN
24
23j45V46
67j47V30
4120
5910
2
180
1500
3702
10,912
SO4
LAC TUY
25 26
23j51V15 23j56V22
67j24V24 67j35V23
4250 4010
5852 5852
1 1
200 180
1500 1500
636 6572
24,433 131,024
SO4 SO4
AC3
27
23j59V16
67j41V33
3950
5910
1
150
1500
2491
25,150
CPR
28
23j57V25
67j47V28
3950
5204
1
150
1500
6589
IMI
29
24j11V06
68j47V00
2949
3995
10
40
2000
PUN
30
24j35V38
69j00V00
2945
6739
10
150
AC4
31
25j00V16
68j37V15
3665
6739
2
LAZ
32
25j04V18
68j30V33
4250
5697
PAJ
33
25j08V33
68j48V50
3537
GOR
34
25j24V07
68j40V27
IGN
35
25j29V39
AZU
36
AMA
0.03
134
0.03
9
1.9 13.4
1.9 13.4
1.6
1.6
306 245
16.2 2.9
2.2 2.9
SO4 – Ca
476
46
2.5
221,804
Ca
137
27
0.9
565
90,008
Ca
189
2000
196
233,367
Ca
4263
180
1630
851
341,759
SO4
656
19.5
1
180
1630
3965
158,854
SO4
393
7.7
5488
5
115
1350
11,728
246,674
Ca
1984
104
1.4
3950
5467
1
140
1000
2846
296,737
SO4
324
27
0.5
68j37V09
4250
5100
2
140
1000
4043
97,091
SO4
25j28V36
68j48V30
3580
5488
3
120
1100
548
323,473
Ca
214
37
25j32V56
68j49V59
3558
5488
2
120
1100
7656
196,672
Ca
863
23
0.04
AGI
38
25j48V06
68j53V24
3320
5058
2
100
1100
177,044
334,882
Ca
589
71
0.0002
ISL
39
25j44V50
68j38V00
3950
5761
0
130
1000
6229
329,693
SO4
858
152
27.5
37.5
9.8 250
0.16 0.1
2
3.5
0.7
0.002
3.3
0.02
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
LAGUNA CHIVATO MUERTO SALAR DE AGUAS CALIENTES 2 LAGUNA LEJIA LAGUNA MISCANTI LAGUNA ˜ IQUE MIN SALAR DE LACO LAGUNA TUYAJTO SALAR DE AGUAS CALIENTES 3 SALAR DE CAPUR SALAR DE IMILAC SALAR DE PUNTA NEGRA SALAR DE AGUAS CALIENTES 4 LAGUNA DE LA AZUFRERA SALAR DE PAJONALES SALAR DE GORBEA SALAR IGNORADO SALAR DE LA AZUFRERA SALAR DE AGUA AMARGA SALAR DE AGUILAR SALAR DE LA ISLA
2.4
253
(continued on next page)
254
Table 1 (continued) Evap (mm/ year)
TDS max (mg/l)
Evol. path
0
140
1000
8907
333,942
SO4
676
40
1.2
5664 5012
2 2
130 100
1000 1100
8277 1677
129,707 318,744
Ca SO4 – Ca
867 293
29 6.7
0.4 0
3494
5373
3
120
1100
3288
20,430
SO4
400
69j08V51
3370
6127
4
125
1200
85
326,745
26j20V12
68j44V49
4150
5857
2
140
1000
2908
82,327
46
26j18V09
68j36V24
4250
6080
2
140
1000
2134
JIL
47
26j23V10
68j39V41
4150
6127
2
140
1000
BAY
48
26j23V09
68j35V02
4250
6127
2
140
WHE
49
26j41V35
68j37V32
4220
6146
1
ESC
50
26j38V51
68j31V21
4353
5918
LAV MAR
51 52
26j52V39 26j55V25
68j27V49 69j04V31
4350 3760
FRA
53
27j27V47
69j13V03
4110
Code
No. (Fig. 1)
Latitude (South)
Longitude (West)
Altitude Salar (m)
Alt max (m)
SALAR DE LAS PARINAS SALAR GRANDE SALAR DE INFIELES SALAR DE LA LAGUNA SALAR DE PEDERNALES SALAR DE PIEDRA PARADA LAGUNAS BRAVAS LAGUNAS DEL JILGUERO LAGUNA DEL BAYO SALAR DE WHEELWRIGHT LAGUNA ESCONDIDA LAGUNA VERDE SALAR DE MARICUNGA LAGUNA DEL NEGRO FRANCISCO
PAR
40
25j49V33
68j30V17
3987
5761
GRA INF
41 42
25j59V52 25j58V32
68j42V00 69j03V26
3950 3520
LGN
43
26j12V29
68j58V29
PED
44
26j13V42
PIE
45
BRA
Temperature mean (jC)
TDS min (mg/l)
SO4 – Ca
Area basin (km2)
Area Salar (km2)
0.55
Area water (km2)
0.4
3620
335
0.6
SO4
388
28
0.2
121,494
SO4
545
10
2925
19,592
SO4
119
3.4
1.5
1000
2432
4785
SO4
221
1.2
1.2
140
1000
1466
158,739
Ca
466
6.3
4
1
140
1000
2318
32,315
SO4
194
3.8
3.5
6893 6893
1 2
170 120
1000 1100
935 144
174,581 331,453
SO4 SO4 – Ca
6052
1
200
1000
130
327,885
SO4
1075 3045 933
15 145 24.8
10
15 6 24.8
Climatic data are taken from the Balance Hidrico de Chile (DGA, 1987). The potential evaporation (Evap) measured in class A evaporimeters has been corrected by a factor 0.65. Alt max is the highest elevation of the drainage basin. TDS min and max are the lowest and the highest salt contents of the waters and brines, respectively, found in each basin. Basin area includes drainage area and salar area. Salar area includes lakes, ponds and dry surfaces (crusts, playas).
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
Prec (mm/ year)
Basin
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293 Table 2 Minerals cited in the text and their formula Amorpous silica Antarcticite Bischofite Calcite Gypsum Halite Mirabilite Natron Thenardite Trona Ulexite
SiO2 CaCl26H2O MgCl26H2O CaCO3 CaSO42H2O NaCl Na2SO410H2O Na2CO310H2O Na2SO4 Na2CO3NaHCO32H2O NaCaB5O98H2O
ranges and basins alternate in the orogenic belt of the Central Andes (Fig. 1). The main morphostructural units are from West to East: (1) the Coast Range; (2) the Central Depression; (3) the Precordillera; (4) the Pre-Andean Depression; (5) the Western Cordillera; (6) the Altiplano; (7) the Eastern Cordillera. The Coast Range is a moderately elevated mountain chain (1500 m as an average) steeply dipping westward in the Pacific Ocean. Marine sedimentary rocks and volcanic andesitic rocks of Mesozoic age make up most of the range. The Central Depression (or Central Valley), at an altitude of 800 to 1400 m, is a thick fill of detrital and lacustrine sediments (gravel, sand, silt, clay) of middle Tertiary to Holocene age, often covered with salt crusts remnants of ancient salars. Roughly between 22j and 25jS latitude, the Central Depression is known as the Atacama Desert. The famous Chilean nitrate deposits are located in this depression where they impregnate many surficial formations (Ericksen, 1981). On the east margin of the valley, large alluvial fan systems make the transition with the Precordillera Range. The Domeyko Range, at an elevation of 3500– 4000 m, is the most conspicuous part of the Precordillera because it is clearly separated from the Eastern Cordillera by the Pre-Andean Depression. The Precordillera Range is made up chiefly of Paleozoic and Mesozoic sedimen-
255
tary, metamorphic and igneous rocks. The large copper ore deposits of northern Chile are located in the Precordillera. The Domeyko Range also contains continental evaporite formations on its eastern fringe. The Pre-Andean Depression, also called the salar de Atacama Depression (not to be confused with the Atacama Desert), is a large intramontane basin, at 2500 m of elevation, filled with Tertiary to Holocene clastic and evaporitic sediments of continental origin. It contains the largest evaporitic basin in Chile: the salar de Atacama ([13], 3000 km2). Another important salar, Punta Negra ([30], 250 km2), occupies the southern edge of the basin. A prominent ridge of anticlinal folds and domes of gypsum and rock salt borders the salar de Atacama on its western fringe: the Cordillera de la Sal. It corresponds to an ancient salar of Tertiary age deformed as a consequence of tectonic events during the Cenozoic (Dingman, 1962, 1967). The Western Cordillera, of Miocene to Holocene age, is an elevated plateau (4000 m) consisting of rhyolitic ignimbrites and covered by numerous andesitic stratovolcanoes reaching 6500 m in elevation. Many of these volcanoes contain native sulfur deposits, some of them of economic interest. The volcanoes often delineate interior drainage basins occupied by saline lakes and salt crusts. Most of the salars included in this study belong to this unit. The Bolivian Altiplano is a major intramontane basin separating the Western Cordillera from the Eastern Cordillera above 22jS latitude. Its central trough is occupied by the giant salar de Uyuni (10,000 km2). The volcanic rocks of the Western Cordillera are in contrast to the sedimentary rocks that dominate in all other units, from the Coast Range to the Precordillera.
4. Climate Climatic features of northern Chile are summarized in Table 3 (DGA, 1987). The potential evaporation
Table 3 Climatic features of northern Chile
Precipitation Potential Evaporation Mean temperature
Coast Range
Central Valley
Precordillera
Pre-Andean Depression
Western Cordillera
< 10 1000 15 – 18
< 10 1500 18 – 20
10 – 25 1500 8 – 12
25 – 50 1500 12
50 – 300 600 – 1200 0
Precipitation and evaporation are in millimeters/year and temperature in jC.
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F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
given in Tables 1 and 3 is derived from measurements in class A evaporimeters (DGA, 1987). Pan evaporation is usually greater than the actual evaporation from lake surface estimated by other methods (hydrologic balance, energy budget). This is due to the differences in size, heat distribution and water movement between large lakes and small and shallow evaporation pans. An empirical pan coefficient is generally applied to correct evaporimeter measurements. The many weaknesses of this approach has been discussed by Miller (1977, p. 266). Unfortunately, at most places, they are the only available data. Reasonable values for pan coefficients may be extrapolated from studies by Jones (1965), Van Denburgh (1975), Yonts et al. (1973), Smith (1979) and Stanhill (2002). Most pan coefficients range from 0.6 to 0.8. In northern Chile, Grilli et al. (1986) selected a single 0.65 correction term, which will be used in this study. A salient feature of the climate is the increase in precipitation from west to east across this region. The coast range and the Central Depression are among the driest deserts of the world, in contrast with the semiarid Western Cordillera where precipitation reaches 300 mm/year. In the whole area, however, the potential evaporation greatly exceeds the precipitation, which is a basic condition for the establishment of evaporitic environments (Hardie and Eugster, 1970; Eugster and Hardie, 1978). In the northern area, 10% to 20% of the precipitation between 4000 and 5000 m occur as snows. In contrast, in the southern area, snow constitutes as much as 50% to 80% of the precipitation at the same elevation (Vuille, 1996). The mean annual temperature decreases from north to south, which allows the snow blanket to persist for a longer time in the southern basins. The dominant wind direction is from ocean to cordillera. Thus, sea salts and saline dust from the Central Valley are transported eastward over the Highlands by the wind.
5. Main outlines of Chilean salars 5.1. Salar morphology A comprehensive description and many photographs of the surficial features of active and fossil
salars, from the Coast Range to the Highlands, is given by Stoertz and Ericksen (1974). Active Chilean salars of the Western Cordillera resemble those described in the Bolivian Altiplano by Risacher and Fritz (1991b). Three main types are found in northern Chile as follows. (1) Saline lakes are permanent bodies of saline waters. The water area occupies the whole salar area (see Table 1). They range in depth from some decimeters to some meters. Mirabilite often precipitates in winter because its solubility is markedly reduced by low temperatures. (2) Playas are the most common type of salar in the Western Cordillera of northern Chile, as they are in the Bolivian Altiplano. Shallow ( < 20 cm) and ephemeral saline pools occupy variable areas of wide mudflats across the central depression of the basin. Some meters below the playa surface, a confined aquifer is filled with an interstitial brine, whose capillary draw precipitates a variety of salts within the sediments (gypsum, mirabilite, halite, ulexite). (3) Salt crusts of two kinds are found in this region. Active salt crusts are presently forming from evaporating brines that have reached saturation with respect to gypsum, mirabilite or halite. These crusts are of low thickness, rarely above 50 cm, and are found in the semiarid Western Cordillera. Fossil salt crusts are located in more arid areas, from the Coast Range to the western fringe of the Western Cordillera. These crusts are of similar mineralogy to the active crusts, but are much thicker. The halite crust of salar de Atacama is several hundred meters thick (Bevacqua, 1991; Bobst et al., 2001). However, fossil crusts are more commonly of metric to decametric size. A 10-mthick halite crust of Neogene age was described in the Coast Range by Chong et al. (1999). Deep dissolution pits (halite, Pajonales [33]), outcrops above the salar level (thenardite, Gorbea [34]), tilted layers (gypsum, Pedernales [44]) are evidence of fossil, nonactive, salt crusts. Many salars are combinations of these types, especially salt crusts and playas. Generally, salars in the Coast Range, in the Central Valley and in the Precordillera are fossil salt crusts, while those in the Western Cordillera are predominantly playa lakes. Mixed salars are found in the Central Depression and in the western fringe of the Western Cordillera,
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close to the Precordillera. In this study, we have incorporated the largest salar of the Central Valley, salar de Pintados, because its underground waters are actively fed by inflows originating in the Western Cordillera. However, the salar itself is a large fossil salt crust.
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et al., 1984; Hastenrath and Kutzbach, 1985; Sylvestre et al., 1999; Baker et al., 2001; Fornari et al., 2001). It is likely that the neighbouring Chilean basins were also filled with saline lakes during the same period. Fossil salt crusts stem from the desiccation of such old saline lakes.
5.2. Inflows 6. Inputs Four types of inflows feed Chilean salars are as follows: (1) Focused discharge through springs from the drainage area with an observable flow rate generally between 0.1 and 100 l/s. Many springs emerge far above the salar surface. (2) Diffuse seeps percolate slowly to the salar surface. Their flow rate is not noticeable, but their lateral extension may reach tens or even hundreds of meters along the salar shore. Total discharge through diffuse seeps may be significant but is impossible to quantify. The slow flow rate may induce an incipient evaporation of the waters at the seepage face. (3) Rivers are generally of short length and low discharge, generally below 500 l/s. They often infiltrate in colluvium before reaching the salars, where they eventually discharge through springs and seeps. (4) Underground waters fill aquifers in the drainage basins. They discharge in the central trough of the basins through springs and seeps. It was often possible to sample these waters in exploration and exploitation wells drilled by mining companies. 5.3. Former lakes Since Miocene time, the climate have remained arid or semiarid in northern Chile (Ko¨tt et al., 1995; Pueyo et al., 2001). Many basins have been occupied by saline lakes (Stoertz and Ericksen, 1974; Geyh et al., 1999). Late Neogene lacustrine formations are known in the Central Valley (Sa´ez et al., 1999; May et al., 1999) and in the Lauca Basin in the Western Cordillera (Gaupp et al., 1999). Grosjean (1994) and Grosjean et al. (1995) described a late Pleistocene lacustrine phase in Laguna Lejia [22]. Several large and deep saline lakes have repeatedly covered the Bolivian Altiplano during Pleistocene time (Servant and Fontes, 1978; Lavenu
6.1. Atmospheric contribution Two fresh snows and seven late-lying snows were sampled and analysed for major and minor components. The meltwater from one old snow has a salt content as high as 23 mg/l due essentially to Ca and SO4, while the salt content of the six others ranges from 1.0 to 6.6 mg/l (mean: 3.6 mg/l). In comparison, meltwaters from fresh snows have salt contents of 3.1 and 3.5 mg/l (mean: 3.3 mg/l). Therefore, six out of seven old snows did not receive significant dry fallout after their deposition, which suggests that these snows were probably sampled shortly after their fall. Only one late-lying snow was markedly contaminated by eolian gypsum. It is therefore possible to obtain a broad estimation of the flux of each component brought by snowfalls in the basins by multiplying the volume of precipitation over each basin by the average concentration of the meltwaters. 6.2. Surface and groundwater inputs On the other hand, the flux of components entering permanent lakes through springs, seeps, rivers and underground discharges can also be estimated. The hydrologic balance of a saline lake without surface or underground outflow is given by (Hutchinson, 1957; Langbein, 1961): Vinflows ¼ S ðE PÞ
ð1Þ
where Vinflows = basin discharge to the lake (m3/year), S = surface of the lake (m2), E = actual evaporation (m/ year) and P = precipitation (m/year). Water loss from Chilean saline lakes through infiltration is generally small compared to basin discharge (see discussion below on salt balance in
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Section 8.2). The flux of components brought in the lake by all inflow waters is obtained by multiplying their average concentration by the volume of inflows. The uncertainty on the basin discharge depends on that of the climatic parameters E and P and on the unweighted average concentration of inflow waters, except when the lake is obviously fed by one major input. Figs. 2 and 3 compare the fluxes of 11 components brought by inflows to 12 permanent lakes to their fluxes deposited by snows on their drainage basins. Lakes are ordered according to the increasing average salt content of their inflows. Two groups can be
distinguished: lakes fed by dilute inflows whose salt content is lower than 500 mg/l (CHR [2], HEL [15], COT [1], LGU [6]) and lakes fed by inflows of higher salt content: 500 –6500 mg/l (MIS [23], LEJ [22], FRA [53], BRA [46], ESC [50], LAV [51], MIN [24], TUY [26]). These two groups of waters differ also significantly in relative ionic concentrations. 6.3. Lakes fed by dilute inflows 6.3.1. Major components in dilute inflows The flux of major components brought by dilute inflows in lakes CHR [2], HEL [15], COT [1], LGU
Fig. 2. Fluxes of major components deposited by snows on the drainage basins (stars) of 12 permanent lakes compared to the fluxes brought into the lakes by inflow waters (squares). Focusing our attention on the SO4 diagram, we can see that the flux of SO4 brought by dilute inflows is only slightly more elevated than that brought by snows on the drainage basins (for LGU [6], the proportion is even reversed). In contrast, concentrated inflows bring in the lakes significantly more SO4 than snows. The other major components show similar behavior although it is not so easily noticeable because of the vertical logarithmic scale.
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[6], are only 1 to 14 times higher than those deposited by the snows (Fig. 2). In LGU basin, all fluxes are even slightly higher in snows than in inflow waters. This means that for LGU, as well as for other lakes fed by dilute waters, atmospheric deposition could represent a major, or the major, contribution to the lakes. This seems particularly clear for SO4 in very dilute waters, which is probably derived from eolian sulfates (gypsum and thenardite) and oxidized elemental sulfur eroded from the top of the volcanoes. Dissolved components in very dilute groundwaters stem from rock alteration and atmospheric inputs that have been brought from the land surface to the groundwater by percolating waters. The flux of atmo-
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pheric inputs deposited on the land surface should then be lower than the flux brought from the basin by inflow waters. A lower flux from the basin could be due to the removal of atmospheric components in soils. We must also lay stress on the large uncertainty on flux estimates. 6.3.2. Minor components in dilute inflows With the only exception of As in Laguna Helada (HEL [15]), the flux of minor components (Br, I, NO3, As) deposited by the snows are 1 to 70 times higher than those brought into the lakes by dilute inflows (Fig. 3: CHR [2], HEL [15], COT [1], LGU [6]), which suggests that most of the minor compo-
Fig. 3. Fluxes of minor components deposited by snows on the drainage basins (stars) of 12 permanent lakes compared to the fluxes brought into the lakes by inflow waters (squares). Except for LGU, most of Si is brought in the lakes by inflow waters. In contrast, the fluxes of Br, I, NO3 and As deposited by the snows are higher than those brought in the lakes by dilute inflows. Not so for concentrated inflows, which bring in the lakes more minor components than snows, with the exception of NO3.
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nents in dilute waters are probably brought by precipitation. Br and I may be derived from sea salts, and NO3 and I may be derived from surficial formations of the Central Valley. The removal of atmospheric components in soils and plants could be very efficient for minor components such as NO3 and I. Mineral deposits and geothermal areas are the primary sources of arsenic in arid regions (Smedley and Kinniburgh, 2002). Owing to the high aridity, As is not leached from the surficial formations and constitutes a noticeable component of desert dust, which makes it a major health hazard in northern Chile (Smith et al., 1998; Queirolo et al., 2000a,b). Its concentration in meltwaters ranges from 0.3 to 68 Ag/l (mean: 18 Ag/l), compared to the world baseline concentration in rain and snow of 0.01 to 0.5 Ag/l (Smedley and Kinniburgh, 2002). In addition, copper smelting plants in northern Chile permanently release large quantities of As in the atmosphere (Gidhagen et al., 2002). As for silica, in all basins but LGU [6], the flux brought by inflow waters is markedly higher (10 to 75 times) than the flux deposited by the snows. Silica is likely to originate from the dissolution of volcanic glass. 6.4. Lakes fed by concentrated inflows Except for NO3, the fluxes of all components deposited by the snows are small in comparison with those brought by concentrated inflows: approximately 3 to 800 times lower. NO3 seems to be the only component, among those studied, that originates in all waters mostly from atmospheric deposition. Br, I, Si and As in these waters must come from some additional source. In the above discussion, we considered only the flux of components brought in the basins by snowfall. However, during the much longer dry season, dry fallout is also deposited on the drainage basins. Therefore, the fluxes of all components brought by atmospheric inputs are very likely to be higher than those estimated only from the snows. This strengthens the conclusion that the lowest the concentration of an inflow water is, the higher is the contribution of atmospheric salts to its salt load. Atmospheric salt is the main source of dissolved components in very dilute waters. As concentration increases, other inputs become progressively more significant.
7. Origin of salts in inflow waters 7.1. Salt content of inflow waters The most conspicuous feature of inflow waters in closed basins of northern Chile is their very broad range of concentration: from 47 to 79,740 mg/l with an average of 3230 mg/l. Northern inflows are globally less saline (average = 2.4 g/l) than southern inflows (average = 4.9 g/l). In comparison, White et al. (1963, 1980) reported salt contents of underground waters in andesitic to rhyolitic aquifers of North America ranging from 70 to 473 mg/l. Alteration modeling of a standard andesite (code Kindisp, Made et al., 1994) leads to a maximum concentration of 340 mg/l of dissolved solids. Thus, the alteration of volcanic rocks does not seem to be the main controlling factor of the high concentration of many inflow waters. The concentration of most components in inflow waters increases with the total salt content (Figs. 4 and 5). The correlation line between component concentration and salt content shows similar slopes for inflows and lakes (Na, Cl, Li>1 and K, Mg, SO4, Ca < 1). The concentration increase in lake waters is easily explained by evaporative concentration, a mechanism that cannot significantly account for most inflow waters. The rate of evaporation of groundwater is much lower than in an open water and decreases drastically with depth (Ripple et al., 1972; Hellwig, 1973; Miller, 1977, p. 304; Yechieli and Wood, 2002). One may reasonably neglect evaporation of deep groundwater within the unsaturated zone. However, shallow groundwater ( < 3 to 5 m below surface) is subject to interstitial evaporation, especially when the capillary fringe reaches the surface. A conspicuous feature of this process is the deposition of efflorescent salts on soils and plants around the salars. Therefore, evaporative concentration may increase the salt content of shallow groundwater at the margin of the salars. However, many deep waters and springs with high discharges also have elevated salt content that cannot be related to direct evaporative concentration. The simplest hypothesis is that dilute waters have mixed with underground brines of composition similar to the present lake brines. The proportion of most components of the deep brines would be nearly conserved by dilution with freshwaters, which could explain the similar slopes observed in Fig. 4.
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Fig. 4. Concentration of major components in inflow waters and lake brines of Chilean salars as a function of total dissolved solids (TDS). Note the overall correspondence between the slopes of the correlation lines of inflows and lakes.
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Fig. 5. Concentration of minor components in inflow waters and lake brines of Chilean salars as a function of total dissolved solids (TDS).
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7.2. Mixing with underground brines Sodium + calcium ( 2) + magnesium ( 2) is plotted against chloride in Fig. 6. Dilute inflows (Cl < 10 mmol/l) present a large scatter in values, while saline inflows (Cl>10 mmol/l) are gathered along the isoconcentration line Cl = Na + 2Ca + 2Mg. Rock alteration is likely to be the main control of the composition of dilute inflows. North American waters from andesitic rocks have compositions similar to the dilute waters. This diagram is alike to those representing any common evaporative system as illustrated in the inset diagram showing a similar relation for evaporated lake waters. However, all waters represented in the main diagram are inflow waters: groundwaters, springs, seeps and rivers. Only seeps and shallow groundwaters may have undergone some evaporation just before they emerge. Nevertheless, seeps and springs present the same behavior, which suggests that the evaporative trend shown in the main diagram could have been inherited from ancient evaporated brines buried below the volcanic formations. Alternatively, the shift to the isoconcentration
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line could be due to the dissolution of a mixture of sodium chloride with calcium and/or magnesium chloride in dilute underground waters. Ancient evaporites containing halite and gypsum are present below the volcanic formations. However, calcium and magnesium chloride are very soluble salts (antarcticite: 1650 g/kg [H2O] and bischofite: 1170 g/kg [H2O]) known to form either in extremely cold environments or in very concentrated brines. They are unlikely to be present in depth at relatively elevated temperature. A more probable explanation is the mixing of dilute waters with Na –Ca – Mg/Cl underground brines similar to those present in saline lakes as shown in the inset diagram in Fig. 6. 7.3. Mixing with subsurface brines Mixing of underground water with brines at the margin of a salar follows the Ghyben – Herzberg principle, which defines the geometry of a freshwater lens of low density lying on a large saltwater body (Heath, 1983). The typical case is a fresh groundwater lens in an island. The Ghyben– Herzberg principle has
Fig. 6. Sodium + calcium ( 2) + magnesium ( 2) vs. chloride concentration for inflow waters (main diagram) and for lake brines (inset diagram) of Chilean salars. North American waters from andesitic rocks are included for comparison. Both diagrams show the same overall pattern, which suggests that inflow waters stem from the mixing of freshwaters with brines similar to those present in the salars.
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received little attention in the case of playas and saline lakes (Yechieli and Wood, 2002). It is difficult to apply to salars in northern Chile. Firstly, the overall pattern is inverted: brine lenses overlie freshwater bodies. This is a basically unstable situation: the denser brine lens tends to move downward. The Ghyben – Herzberg mixing zone slopes away from a saline lake (see Yechieli and Wood, 2002), in contrast with a freshwater lens overlying a saline layer (see Heath, 1983, p. 68). It is not easy to extrapolate very much in depth the interface between fresh and saltwater below a saline lake. The geometry of the lens can be reasonably described only close to the surface, near the zone of discharge. Secondly, the Ghyben – Herzberg principle applies only to an isotropic aquifer, which is not the case of volcanic formations with fracture porosity. Water and brine may be forced to move along unpredictable pathways. Thirdly, thermal heating below the basin may considerably disturb the interface between the water bodies. Anyway, brine mixing occurs at the margin of the salars and could be partially responsible of the elevated salt content of the seeps. However, many springs emerging far above the
salar surface also have high salt content. Thus, mixing at the margins is not likely to be the main process of the high concentration of inflow waters. 7.4. Geochemistry of bromide Chloride-to-bromide ratios shown in Fig. 7 also suggest that the composition of saline inflow waters (Cl>10 mmol/l) results from the mixing of snow meltwaters with brines similar to those found in the present lakes. Two mixing curves between average snow meltwater and lake brines are drawn: curve A corresponds to the lake brine of lowest Cl/Br ratio and curve B to that of highest Cl/Br ratio. Any other mixing curve is constrained between these two end member curves. All saline inflows plot in this mixing area. By contrast, several dilute inflows (Cl < 10 mmol/l) plot outside the mixing area. Br content of halite associated to Tertiary gypsum diapirs in the neighbouring Bolivian Altiplano is around 2 ppm (Risacher and Fritz, 2000). Similar low values (1– 2 ppm) have been found by Pueyo et al. (2001) in neogene salt rocks of northern Chile. The Cl/Br ratio
Fig. 7. Cl/Br ratio as a function of Cl for snows, inflow waters and lake brines. Two end member mixing curves between snow meltwater and lake brines are drawn: Curve A corresponds to the lowest Cl/Br ratio (alc-9) and curve B corresponds to the highest ratio (isl-12). All other possible mixing curves are located between these two curves. Curve C shows the Cl/Br ratio of the solution produced by dissolving halite in snow meltwater. Curve C is well outside the field of saline inflows.
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obtained by dissolving such low Br content halite in pure water is figured by an horizontal dashed line far above all inflow points. Curve C represents the Cl/Br ratio of the solution obtained by dissolving halite in average snow meltwater. The halite endpoint may be put anywhere on the horizontal dashed line for Cl values between 355 and 46,500 mg/l, which is the highest Cl content measured in saline inflows. The dissolution curve C if well above all saline inflow points. Thus, the high salt content of inflows cannot stem from the dissolution of ancient salt rock overlapped by volcanic formations. 7.5. Stable isotopes The d2H and d18O values for 89 inflow waters, 6 lakes (Chaffaut, 1998) and 26 snows (Pen˜a, 1989) are shown in Fig. 8. The meteoric water line for northern Chile is taken from Aravena et al. (1999). The lakes were sampled in the central part of the studied area ( f 23jS latitude) and the snows in the northernmost sector ( f 18j). Most of the snows show a marked displacement to the right of the meteoric line. According to Pen˜a (1989), the shift is due to sublimation and evaporation of the snow blanket during the day. Night
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condensation does not counterbalance the evaporation in this arid region. Moser and Stichler (1980) carried out experimental evaporation of snow samples which led to a similar shift due to d2H and d18O increase. Inflow waters are clearly distributed along a line joining the lightest isotopic values to the heavily enriched ones observed in the lakes. The isotopic values of inflow waters follow geographic and hydrologic trends. (1) Northern inflows ( < 24j latitude) have lower isotopic values than southern inflows; only two northern inflows have high isotopic values similar to those found in the south. (2) Seeps are more enriched in d2H and d18O. The shift of inflow waters to the right of the meteoric line is due to the conjunction of several processes reflected by the two correlations. Snow prevails over rain and remains for a longer time in the south than in the north of the study area (see Section 4). As a result, the evaporation of the snows and their related isotopic enrichment observed in the north by Pen˜a (1989) is likely to be significantly enhanced in the south. Seep waters are subjected to evaporation by capillary draw before they emerge and by direct evaporation as they percolate very slowly to the land surface. There is no surprise if these waters have
Fig. 8. 18O and 2H contents in snows, inflow waters, and some lake brines of Chilean salars. The isotopic composition of the most concentrated inflow of salar de Surire [4] that results from mixing of groundwater with brine is shown. The size of lake symbols (open circles) is not related to their salt content.
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Fig. 9. Ionic activity products of calcite, amorphous silica and gypsum as a function of total dissolved solids (TDS). Square brackets refer to activities. Horizontal lines show the solubility product at 0 and 25 jC. A water is supersaturated with respect to a phase when its representative point plots above the solubility product line. Below the line, the water is undersaturated with respect to the phase. Many inflow waters are oversaturated, or close to saturation, with respect to calcite and amorphous silica. These two minerals seem to exert a main control on Si, Ca and carbonates concentration of most inflow waters. As TDS increases, inflow waters tend to reach saturation with respect to gypsum. Eventually, gypsum is saturated, or very close to saturation, in 3 springs and 10 seeps, which suggests gypsum dissolution, at least in spring waters.
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high d2H and d18O values. In contrast, underground waters, springs and rivers (of short length in the study area) do not undergo significant evaporation before sampling, but also present a significative isotopic enrichment, particularly in the south. The higher salt content of many of the southern waters associated with heavier isotopic contents suggests rather a mixing process than an evaporative concentration. Dilute and concentrated endpoints are snow meltwaters and brines, respectively, similar to those found in present saline lakes. The relation between salt content and isotopic enrichment suggested in Fig. 8 is partially obscured by the respective salt contents of the two endpoints. For example, the water resulting from the mixing of a little volume of hypersaline brine with a large volume of dilute water has an isotopic composition close to that of the dilute water. In contrast, the salt content of the mixed water may be markedly higher than that of the dilute water. The volume of saline water is negligible, but not the mass of dissolved salts. A typical example is that of the most saline inflow (17.7 g/l) to the salar de Surire [4]. The spring emerges within the salar after passing through a thick section of brine-filled sediments. The high salt content
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is obviously not due to evaporative concentration of the spring water, but to the mixing with a little volume of hypersaline brine, which is confirmed by the d2H and d18O values plotting almost on the meteoric water line. 7.6. Mineral dissolution Fig. 9 shows the ionic activity product of three important minerals involved in the salt load of inflow waters and in their evaporative evolution: calcite, gypsum and amorphous silica. Solute and water activities are calculated with the ion interaction model (Pitzer, 1979; Harvie et al., 1984). H4SiO4 activity and amorphous silica solubility are taken from Marshall (1980), Marshall and Warakomski (1980) and Marshall and Chen (1982). The software used is EQL/ EVP (Risacher and Clement, 2001 and complete database references herein). Most of the inflow waters are saturated, or close to saturation, with respect to calcite and amorphous silica. Therefore, the solubility of amorphous silica controls the concentration of silica in inflow waters. Calcite is very likely an alteration mineral that controls, at least partially, the concentration of Ca
Fig. 10. Temperature (jC) of springs and underground waters and mean annual air temperature as a function of elevation above sea level. Waters that plot above the curve of mean annual air temperature + 5 jC present a thermal influence.
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and HCO3 in inflow waters. Most inflow waters are undersaturated with respect to gypsum. However, as concentration increases, gypsum becomes close to saturation in 10 seeps and 3 springs. Concerning the seeps, it is possible that their Ca and SO4 concentration has increased by evaporation. However, this cannot be the case for springs. For these waters, the dissolution of gypsum is the simplest way to reach saturation. There are no other common minerals that
could supply sufficient quantities of Ca and SO4 to reach saturation with gypsum. 7.7. Thermal waters Northern Chile is a region of high heat flow due to recently active volcanism (Springer and Fo¨rster, 1998). Geothermal systems, such as that of El Tatio (outside the study area), have been considered for
Fig. 11. Lithium and boron content of springs and underground waters [in wt.% of total dissolved solids (TDS)] as a function of temperature. No correlation with temperature is observed.
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energy production (Lahsen, 1988). Thermal influence can only be estimated in springs and underground waters. The temperature of seeps and rivers is controlled by the atmosphere. The temperature of springs and underground waters is compared in Fig. 10 to the mean annual air temperature (Grilli et al., 1986; DGA, 1987). If we reasonably assume that a water whose temperature is at least 5 jC higher than the mean annual air temperature presents a thermal influence, then thermal heating has affected 78% of all sampled springs and underground waters. However, the thermal influence is moderate: only 10 waters have temperatures between 40 and 53 jC, which are the highest values observed in the study area. No correlation between temperature and solute concentration
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has been found. Fig. 11 shows the concentration of Li and B (in wt.% of the salt content) as a function of water temperature. Li and B are generally enriched in thermal waters (Berthold and Baker, 1976; White et al., 1976). Nevertheless, in northern Chile, such an enrichment is not observed. Therefore, hydrothermal alteration does not seem to be a major process of solute acquisition. Fig. 12 shows the spatial distribution of inflows differentiated according to the difference between their temperature and that of the atmosphere (Twater –Tatm; seeps and rivers are not considered). In the northern and central area (18 – 24jS latitude), both thermal and nonthermal waters are intermingled. In the south (24 – 27jS latitude), however, a clear separation can be observed. Nonthermal springs are located westward close to the boundary between the sedimentary Precordillera and the volcanic Western Cordillera. In contrast, heated inflows are found preferentially eastward well within the volcanic Cordillera. This observation is in good agreement with the increase of heat flow from the Coast Range to the active magmatic arc of the Western Cordillera reported by Springer and Fo¨rster (1998).
8. Origin of deep brines 8.1. Recycling of ancient salars
Fig. 12. Distribution of thermal springs and underground waters in salar basins of northern Chile. Note the exaggerated horizontal scale. In the south, thermal inflows are more abundant eastward within the Cordillera. A similar relation may also be suspected in the northwestern part of the study area.
Volcanic formations in northern Chile range in age from Miocene to Pleistocene. During this entire time span, the climate remained arid or semiarid (Ko¨tt et al., 1995; Pueyo et al., 2001). Many salars must have occupied ancient intravolcanic basins. Subsequently, lava flows and ignimbrite sheets covered these saline deposits. The buried brines may have mixed with dilute meteoric inflows to produce brackish and saline underground waters. However, this straightforward explanation for the origin of deep brines comes up against two difficulties. First, the brines must have been preserved during the hot volcanic events. Elevated temperatures may vaporize the lake waters. Only underground brines protected by overlying sediments could have been preserved. Secondly, these entrapped brines must have been rapidly exhausted as they discharge with fresh inflows. The time of exhaustion depends of the size of the buried salars
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and of their leaching rate by dilute waters. We may intend a rough calculation of the exhaustion time based on present salars. Their wide range of morphological and chemical types make it difficult to define an average salar. On the other hand, extreme values of salar size and water composition may not be very significative. Therefore, we have selected a broad range of values for each parameter that covers a large number of salars of different type: salar surface = 10 to 150 km2; lake surface = 1 to 15 km2; salt content of inflows = 0.5 to 5 g/l; salt content of brines = 20 to 200 g/l; evaporation = 1.33 m/year; precipitation = 0.16 m/year. The only parameter that cannot be easily estimated is the volume of underground brines. We will assume a thickness of 100 m of brinefilled sediments with a porosity of 40%. The mass of dissolved salts in deep brines of a buried salar can be estimated by multiplying the salar surface by the sediment thickness, the porosity and the brine concentration. The flux of dissolved salts brought each year in a present salar is obtained by multiplying the volume of inflows [see above, Eq. (1)] by their salt content. By dividing these two values, we obtain an exhaustion time ranging from 100 years (a small salar rapidly leached, generating concentrated inflow waters) to 1.7 million years (a large salar slowly leached, generating dilute waters). Small- and medium-sized salars have been leached in a short time as compared with the age of volcanic formations of Central Andes (Miocene to Holocene). Only large buried salars could still provide some salts in present inflow waters. 8.2. Brine leakage from present salars Brine leakage from present salars is another source of salts in underground waters. The concentration of a conservative component (Li, Br, Cl before halite saturation) in a saline lake is controlled by the infiltration rate of the lake brine through bottom sediments. Water leakage as a controlling factor of lake chemistry has already been stressed by Carmouze and Pedro (1977), Alderman (1983), Sanford and Wood (1991) and Dutkiewicz et al. (2000). If a lakebed is completely impermeable, then the concentration of a conservative component, and consequently the salt content of the lake, will continuously increase, which in turn reduces the evaporation rate of
the lake brine. A steady state is attained at very high concentration when the equilibrium water vapor pressure exerted by the solution equals the mean relative humidity of the atmosphere (Kinsman, 1976). In contrast, if infiltration occurs through the lakebed, even at a very slow rate, then a steady state is attained when the input flux balances the leakage flux of the conservative component. The difficulty is to determine if the steady state is attained. A recent climatic, geologic or anthropic event may have abruptly modified the lake composition, which could be undetectably reequilibrating through time with the new climatic or geologic conditions. One way is to compare the total mass of the component in the lake to its annual input flux. The ratio between total mass and annual input has the dimension of a time. If this time is substantially lower than the age of the last disturbing event, then it is referred to as the residence time of the component in the lake. A short residence time indicates a high infiltration rate. The total mass of a component in the lake is: Mlake ¼ Slake Hlake Clake
ð2Þ
where Slake is the surface of the lake, Hlake is its mean depth and Clake is the average concentration of the selected conservative component in the lake. The annual input flux is [see Eq. (1)]: Minflows ¼ Slake ðE PÞ Cinflows
ð3Þ
where E is the actual evaporation, P is the precipitation and Cinflows is the average concentration of the conservative component in inflows. The ratio between the total mass in the lake and the annual input flux is: Mlake =Minflows ¼ Hlake =ðE PÞ Clake =Cinflows
ð4Þ
Table 4 presents this ratio for Cl in some characteristic Chilean saline lakes. Most lakes belong to the Alconcha [10] – Lagunilla [6] types with ratios ranging from a few years to some hundreds. Laguna Helada [15] has the highest estimated ratio: 3370 years. The last drastic climatic change occurred 10,000 years ago when large lakes occupying the basins dried up (see Section 5.3). It seems therefore reasonable to assume that all these ratios represent the residence time of Cl and that Andean saline lakes are under steady-state
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Table 4 Rough estimates of residence times and infiltration rates of some typical north Chilean saline lakes Lake surface (km2) Mean lake depth (m) Mean inflow concentration (Cl: mmol/l) Lake concentration (Cl: mmol/l) Residence time [years, Eq. (4)] Infiltration rate [mm/year, Eq. (5)] Loss through infiltration (106 l/year) Basin discharge [106 l/year, Eq. (1)] Percent infiltration relative to basin discharge
HEL [15]
FRA [53]
ALC [10]
MIS [23]
LGU [6]
5.8 0.50 0.64 5050 3370 0.15 0.86 6800 0.01
14.3 0.20 2.0 4810 411 0.5 7.0 16,700 0.04
0.75 0.20 0.80 1030 220 0.90 0.68 880 0.08
13.4 4.5 1.8 34 73 62 830 15,700 5.3
0.15 0.15 21 161 1 150 23 180 13
Cl is the conservative component used for the calculations. Lake levels may change during the year. Cl concentration corresponds to the lake size given in this table. FRA surface is different from that given in Table 1. Actually, FRA is constituted of two separated lakes: one dilute lake that acts as a preconcentration pool and a hypersaline lake which is considered here. MIS mean depth is taken from Valero-Garce´s et al. (1996) and FRA mean depth from Behnke (1987). In both cases, we have halved the maximum depth given in these studies to obtain an approximate mean value inasmuch as the surface of both lakes reported by these authors was similar to that observed during our field trips. All other mean depths were estimated by ourselves. The main uncertainty is the mean Cl concentration of inflows, difficult to estimate. Underground discharges and ephemeral streams during the wet season are not taken into account.
regime as well as most subsurface brines. The salt balance for a conservative component (Cl, Li, Br) is then given by: Vinflows Cinflows ¼ Vleakage Clake
ð5Þ
where C is the average concentration of the component and V is the annual flux of solutions. Eq. (5) allows to estimate leakage rates from saline lakes. Table 4 gives infiltration rates of characteristic Chilean saline lakes. The volume of lake brine lost by infiltration ranges from 0.01% to 12.8% of the basin discharge. Therefore, it is generally justified to neglect the leakage in the water balance of a saline lake although it is the main control of the salt balance of the lake. Almost all dissolved salts entering the lake are lost by leakage. Only low soluble salts (calcite, gypsum, Mg silicates) are removed from the solution and stored in lake sediments. Their mass is very small compared to the total input flux to the lake. 8.3. Dissolution of present salt crusts Most of the incoming solutes are lost by infiltration, which precludes the formation of thick salt crusts. However, some salars, particularly those located close to the Precordillera, contain several meter thick halite crusts filled with an interstitial
brine (Imilac [29], Pajonales [33], Gorbea [34], Agua Amarga [37], Aguilar [38], Pedernales [44]). These thick crusts are quite different from those of centimetric to decimetric size forming by evaporation in other salars (see Section 5.1). Numerous deep brine ponds suggest that the thick crusts are undergoing active dissolution. They do not seem to be presently forming. Under former wetter climate, deep salt lakes occupied the basins (see Section 5.3). An abrupt drying-up due to a climatic change induced the rapid evaporation of the lake and the precipitation of a thick salt crust. The residual lake is often limited to a subsurface, intracrustal brine saturated with respect to halite. The intracrustal brine infiltrates through bottom sediments removing large quantities of dissolved salts. In contrast, inputs to the salar are fresh or brackish waters largely undersaturated with respect to halite, which tends to dilute the intracrustal brine. In order to maintain saturation with respect to halite, the crust must dissolve. As long as the salt crust remains in the salar, the flux of infiltrating salts is higher than the flux of incoming salts, which induces a slow dissolution process responsible for the numerous ponds pitting the crust. A new steady-state regime will be established only when the whole crust has vanished. In summary, ancient salt crusts in present Andean salars are not likely to be in equilibrium with their climatic and hydrologic environment. They are an additional
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source of salts in underground waters close to the Precordillera. 8.4. The salt cycle in Chilean salars Salts do not accumulate in most Andean salars. The salt cycle is only delayed. The delay is the residence time of the dissolved components. Most incoming salts only pass through the saline formations, where they undergo a concentration process and are further recycled in the hydrologic system. Infiltrating brines are not necessarily recycled in the same basin. The intense and complex fracturing of volcanic rocks could favor deep circulation from one basin to the others. This is a continuous mechanism that accounts well for the high salt content of inflow waters over long time spans. Thermal convection is the likely driving force that mixes, moves and heats the underground waters in this area of active volcanism. There is no need of a large salt reservoir below the volcanic structures. Nevertheless, giant salt bodies are present in depth. Alonso et al. (1991) estimate at 10,000 km3 the volume of deep halite in Central Andes. The Bolivian Altiplano is dotted with gypsum diapirs. The Cordillera de la Sal in northern Chile plunges to the east below ignimbrites and lava flows. Gypsum recycling through volcanic activity in the Western Cordillera has been reported by Risacher and Alonso (2001). Presently, only gypsum seems to supply part of Ca and SO4 to the salars. Owing to its high solubility, the top of halite bodies formerly in contact with underground waters could have been leached a long time ago and redeposited in the giant salars of Uyuni and Atacama. Only low soluble gypsum is still in contact with subsurface waters and constitutes a present source of Ca and SO4.
9. Brine evolution 9.1. The model Waters undergoing evaporative concentration precipitate a sequence of minerals in order of increasing solubility. Hardie and Eugster (1970) and Eugster and Hardie (1978) introduced the basic concept of chemical divide to account for the evolution of solute concentration when a water evaporates. In brief sum-
mary, when a mineral precipitates, the concentration of all its components cannot increase simultaneously in the solution because the ionic activity product of the mineral must be kept constant and equal to its solubility product. For example, in the case of a mineral containing one anion and one cation, if the anion activity increases in the solution, then the cation activity must decrease. The solution becomes enriched in some components and depleted in others, according to the precipitated minerals and to the concentration ratio of their components at the start of the precipitation. Two approaches may be used to predict the fate of an inflow water upon evaporation. The quantitative approach calculates step by step the chemical composition of the solution and the sequence and quantities of minerals produced during evaporation. In this study, we use the EQL/EVP code (Risacher and Clement, 2001) based on the ion interaction model (Pitzer, 1979; Harvie et al., 1984). The qualitative approach only intends to predict which pathway a water will follow, without caring about the precise chemical composition of the solution. The main difficulty lies in that many analyses found in the literature, and all of this study, report total alkalinity (Alk) as a routinely measured parameter. The first chemical divide produced by the early precipitation of calcite would apparently necessitate to calculate the respective concentrations of CO32 and HCO3. The first model established by Hardie and Eugster (1970) was grounded on carbonate speciation, neglecting ion association and using the Debye – Hu¨ckel equation for ion activities. The [Ca2 +] vs. [CO32 ] concentration relation is a rather complex expression because [HCO3] and [OH] must also be taken into account. In addition, each analysis must be recomputed so that the water is in equilibrium with the atmospheric pCO2. It is possible to update these calculations by using the ion interaction model and the EQL/EVP code. However, this would be a useless repetition of the quantitative approach. A much simpler and rigorous treatment directly based on the alkalinity concept has been developed by Al-Droubi et al., (1980) for the first chemical divides in which carbonate and silicate species are involved. We will summarize this approach which does not require any computer code (see also Risacher and Fritz, 1991b).
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The total alkalinity of a natural water can be expressed as (Dickson, 1981; Stumm and Morgan, 1996): Alk ¼ 2½CO2 3 þ ½HCO3 þ ½OH þ ½BðOHÞ4
½Hþ
ð6Þ
where square brackets refer to total concentrations in mol/l or mmol/l. This formulation corresponds to the operational definition of alkalinity, but is difficult to handle because all terms are interdependent. By combining Eq. (6) with the electro-neutrality equation, we obtain the alternative definition of alkalinity: Alk ¼ ½Naþ þ ½Kþ þ 2½Ca2þ þ 2½Mg2þ ½Cl 2½SO2 4
ð7Þ
Alkalinity is the difference between the sum of the conjugate cations of the strong bases and the sum of the conjugate anions of the strong acids (Stumm and Morgan, 1996, p. 165). Calcite is the first mineral to precipitate from most natural solutions undergoing evaporative concentration. We may rewrite Eq. (7) as: Alk 2½Ca2þ ¼ ½Naþ þ ½Kþ þ 2½Mg2þ ½Cl 2½SO2 4
ð8Þ
As long as calcite is the only precipitate, the concentration of each component of the right-hand side of Eq. (8) increases linearly with the concentration factor F of the evaporating solution: ½Naþ ¼ F ½Naþ 0
ð9Þ
where F ¼ ðH2 OÞ0 =ðH2 OÞ
ð10Þ
(H2O) is the mole number of water in the solution. Subscript ‘‘0’’ refers to the initial solution at the start of evaporation. Eq. (8) becomes: Alk 2½Ca2þ ¼ F ð½Naþ 0 þ ½Kþ 0 þ 2½Mg2þ 0 ½Cl 0 2½SO2 4 0 Þ
ð11Þ
which leads to: Alk 2½Ca2þ ¼ F ðAlk0 2½Ca2þ 0 Þ
ð12Þ
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If at the start of evaporation Alk0>2[Ca2 +]0, then the right-hand side of Eq. (12) is positive and increases linearly with the concentration factor. Thus, as evaporation proceeds, the difference between alkalinity and calcium concentration will continuously increase on behalf of alkalinity. Conversely, if Alk0 < 2[Ca2 +]0, then calcium concentration will increasingly predominate over alkalinity. Inasmuch as carbonate species are the main contributors to alkalinity in most natural waters, the behavior of alkalinity reflects that of carbonates. Therefore, the first chemical divide of a water composition upon evaporation can be very easily and rigorously determined by examining the simple ratio of the analytical concentrations Alk0/ 2[Ca2 +]0 of the initial inflow water. Mg carbonate or Mg silicates also precipitate at an early stage of water evolution. The effect of coprecipitation of calcite and Mg salts on the evolution of an evaporating water may be determined by rewriting Eq. (7) as: Alk 2½Ca2þ 2½Mg2þ ¼ ½Naþ þ ½Kþ ½Cl 2½SO2 4
ð13Þ
which leads straightforwardly to: Alk 2½Ca2þ 2½Mg2þ ¼ F ðAlk0 2½Ca2þ 0 2½Mg2þ 0 Þ
ð14Þ
If Alk0>2[Ca2 +]0 + 2[Mg2 +]0, then the difference between Alk and 2[Ca2 +] + 2[Mg2 +] is positive and continuously increasing. The evaporating water evolves into an alkaline brine depleted in Ca and Mg and enriched in carbonates. Conversely, if Alk0 < 2[Ca2 +]0 + 2[Mg2 +]0, then the water follows a neutral path depleted in carbonate. If Alk0>2[Ca2 +]0 and Alk0 < 2[Ca2 +] 0 + 2[Mg2 +]0, then the evaporating water first follows the alkaline path as long as calcite is the sole precipitate and reverts to the neutral path when Mg carbonate or Mg silicates reach saturation. It must be emphasized that either Mg carbonate or Mg silicates have the same effect upon the chemical divide. Fig. 13 shows some initial evolutionary paths of north Chilean inflow waters in a diagram Alk vs. 2[Ca]. Waters following the neutral path become progressively enriched in Ca and, in most cases, reach saturation with gypsum (CaSO 42H2O). Gypsum
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Fig. 13. Examples of evaporation paths of north Chilean inflow waters illustrating the qualitative approach (analyses in Risacher et al., 1999: 1 = asc-14; 2 = mic-1; 3 = ped-9; 4 = sur-26; 5 = sur-7). Solid circles are directly plotted from water analyses. Quantitative evaporation paths are computed with the EQL/EVP evaporation program based on the ion interaction model and a rigorous treatment of the aqueous species. At the beginning of evaporation, all waters remain undersaturated with respect to calcite. Ca and Alk behave conservatively and each water follows a straight line path whose equation is 2[Ca]/Alk = 2[Ca]0/Alk0. 2[Ca]0 and Alk0 are the coordinates of the starting points (solid circles 1 to 5). The precipitation of calcite causes an abrupt deviation of the linear evaporation path depending on which side of the iso-concentration line Alk = 2[Ca] the starting point stands. The precipitation of magnesite in path 5 reverses the alkaline path to the neutral path because Alk0 < [Ca]0 + 2[Mg]0 in inflow water 5. The precipitation of Mg salts also occurs in paths 3 and 4, but does not reverse the alkaline paths inasmuch as Alk0>2[Ca]0 + 2[Mg]0 in these waters. The pH of the initial solution may be modified by CO2 exchange with the atmosphere or by CO2 uptake by aquatic plants. The result is a displacement of the calcite appearance point along the initial conservative straight line, which does not change the evolutionary trend.
induces a new divide of the neutral path producing either Ca-rich/SO4-poor waters or Ca-poor/SO4-rich waters, according to their Ca/SO4 ratio at the beginning of gypsum saturation (not at the beginning of evaporation). The whole discussion is summarized in the flow sheet of Fig. 14 which accounts for most natural waters (modified from Hardie and Eugster, 1970). A few minerals play a prominent part in the water evolution. The outcome of the processes is the generation of three main groups of brines: alkaline (Na/ HCO3 –CO3 –Cl), sulfate-rich (Na/SO4 – Cl) and calcium-rich (Na –Ca/Cl) brines. The flow sheet of Fig. 14 defines four pathways: alkaline (I – IA); sulfatealkaline (I – IIA –III); sulfate-neutral (II – III) and calcic (II – IV). Pathways I– IIA –IV and II –(Na/CO3 – Cl) have not been observed in nature.
Actually, Na and Cl are the main components of most brines, but they do not induce a chemical divide before halite saturation, which occurs at high concentration once the major evolutionary trends of the brine have already been defined. Throughout the first crucial steps of evaporative concentration, Na and Cl behave as conservative components. For the sake of brevity, waters, where SO4>Ca, will be referred to as SO4-rich, and those where Ca>SO4 will be referred to as Ca-rich. The composition of a brine is strongly dependent on that of the initial dilute water. A small variation of the ratio of the critical components in the dilute water {Alk0/2[Ca]0, Alk0/(2[Ca]0 + 2[Mg]0)} may cause drastic changes in the final brine composition. In turn, the composition of dilute inflow waters is mainly inherited from rock alteration. Therefore, there is an
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Fig. 14. Flow diagram for evaporative concentration of dilute waters. Path numbers are those of Hardie and Eugster (1970): alkaline (I – IA); sulfate-alkaline (I – IIA – III); sulfate-neutral (II – III) and calcic (II – IV). The dashed line shows a rarely observed direct path from the Mg salts divide to the sulfate-rich brine group, bypassing the gypsum divide. It has no particular incidence for water evolution and is included in our discussion in the sulfate-alkaline path I + IIA + III.
overall relation between lithology and brine composition. Each broad type of lithology is reflected in the composition of the dilute water which is further amplified by evaporative concentration. In a pioneering study, Garrels and Mackenzie (1967) first threw light on the (nonobvious) relation between igneous rocks and alkaline brines. High-purity igneous rocks are almost devoid of anionic components (Cl , SO42 ). The only available anion that can balance the cations derived from the weathering of pure igneous rocks is HCO3, which stems from atmospheric CO2 and is brought by the rains as carbonic acid. However, igneous rocks often contain mineralization or Cl – SO4-rich fluid inclusions which lowers the proportion of HCO3 in the weathering solution. Moreover, sulfide minerals and native sulfur not only
provide SO42 but also acidify the solution through oxidation reactions such as: Sj þ 3=2O2 þ H2 O ! 2Hþ þ SO2 4
ð15Þ
The alkalinity of the solution may be drastically reduced. Therefore, inflow waters draining moderately mineralized igneous rocks tend to follow the sulfate alkaline evaporative pathway, while highly mineralized igneous rocks produce waters following the sulfate-neutral pathway. Extreme cases of very intense acidification leading to acid-sulfate brines have been reported in Australia (McArthur et al., 1991; Long et al., 1992) and Chile (Risacher et al., 2002). Sedimentary rocks are often characterized by high Ca (limestone, marl, dolomite, gypsum). Waters
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in contact with such rocks generally tend to generate calcic brines. 9.2. Discrepancies between model and observed brine evolution The model relating lithology, inflow waters and brine composition sets the geochemical and hydrochemical baseline of evaporative concentration. It assumes direct evaporation of the inflow water as it would occur in an artificial pool protected from any contamination. Actually, several processes may disturb this scenario. The evaporating water often interacts with eolian dust, sediments or salts, which modifies its composition and may change its theoretical evolutionary path. For example, the behavior of potassium is mostly controlled by exchange reactions on clay minerals (Eugster and Jones, 1979). The content of SO4 may be lowered by bacterial reduction or freezing out of Na sulfate salts (Eugster and Hardie, 1978). Gac (1980) carried out evaporation experiments of the main tributary of Lake Chad with and without allowing contact with atmospheric dust. The behavior of Mg was very different in each experiment. In Bolivian Andes, native sulfur is eroded from the top of the volcanoes and deposited on the drainage basins where its oxidation acidifies inflow waters enough to reverse their initial alkaline evolutionary path to the sulfatealkaline or sulfate-neutral path (Risacher and Fritz, 1991b). These interactions are most efficient at an early stage of evaporative evolution, when inflow waters are still dilute. Brackish and saline inflows are much less sensitive to them. Another disturbing mechanism is the mixing of several inflows with different evolutionary paths. Mixing in large basins of complex lithology is a more likely scenario than in small monolithologic basins. Leakage through bottom sediments has also been reported to significantly modify the brine evolution (Sanford and Wood, 1991). All these well-documented processes take place after discharge of inflow waters in the basin, while they undergo evaporative concentration. In this paper, we describe a (apparently) new discrepancy between lithology and evolutionary paths of inflow waters. Several Chilean salars located in an exclusive volcanic environment are fed by inflow waters following the calcic path typical of sedimentary rocks. Their brines do not reflect the geological context.
9.3. Water evolution in Chilean salars 9.3.1. The calcite divide As shown in Fig. 9, most inflow waters are close to saturation, or oversaturated, with respect to calcite. The relations Alk vs. 2Ca and Alk vs. 2Ca + 2Mg for inflow waters are shown in Fig. 15. At the start of evaporation (upper diagram), 238 inflow waters (67%) should follow the initial neutral pathway (II), while 117 (33%) should follow the opposite initial alkaline pathway (I). The precipitation of Mg salts (lower diagram), probably Mg smectites as in Bolivian salars (Badaut and Risacher, 1983), theoretically shifts 65 out of 117 waters from the initial alkaline pathway to the sulfate-alkaline pathway (I – IIA), remaining 52 waters (15%) that should end up as carbonate brines (I – IA). In contrast to the prediction, in the whole study area, only one small brackish pond located in the drainage basin of laguna Helada [15] is actually of alkaline composition (HEL [15] in the lower diagram). All other 226 lakes and ponds contain low-alkaline or neutral waters of the Na/Cl– SO4 or Na – Ca/Cl groups including the main laguna Helada itself. The discrepancy has already been observed and discussed in the neighbouring Bolivian salars and explained by acidification due to the oxidation of eolian sulfur in the drainage basins (Risacher and Fritz, 1991b). It is very likely that the same process occurs in northern Chile, where many volcanoes also contain similar native sulfur deposits. However, the proportion of inflow waters that should end up as carbonate brines in Bolivian Andes is much higher than in northern Chile: 52% and 15%, respectively. Moreover, 6 Bolivian salars out of 30 are actually alkaline. Therefore, an additional process must be involved to explain the lower proportion of alkaline waters and brines in Chilean Andes. One of the most likely is dry fallout of desert dust from the Central Valley on the Chilean Andes. The importance of atmospheric inputs in the composition of very dilute waters has already been emphasized (see Section 6). Gypsum is one of the main components of desert and atmospheric dust. Its dissolution increases the concentration of Ca of the inflow waters without modifying their alkalinity, which may shift the water evolution from the alkaline path to the neutral or to the sulfate-alkaline paths.
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Fig. 15. 2Ca and 2Ca + 2Mg concentration vs. alkalinity for inflow waters in Chilean salars. Lake brines values are also plotted on the 2Ca + 2Mg diagram. To avoid redundancy, waters with Alk < 2Ca are not plotted in the lower diagram. Inset flow diagrams show the evaporative concentration paths followed by the group of waters differentiated by the isoconcentration lines. There is only one lake brine (HEL) in the alkaline-rich field although 52 inflow waters should end up after evaporation in the alkaline-rich field.
9.3.2. The gypsum divide The upper diagram of Fig. 16 shows the Ca and SO4 contents of all inflows and lake waters. Lake waters are differentiated according to their Ca/SO4 ratios: 143 are sulfate-rich and 82 Ca-rich waters. In contrast, inflow waters are differentiated according to their evolutionary path calculated with the EQL/EVP code (Risacher and Clement, 2001). Many inflow waters that end up as sulfate brines have initial concentrations higher in Ca than in SO4. The early
precipitation of calcite reverses this ratio allowing the solution to become sulfate-rich. Two typical evolutionary paths are drawn: one (line A) leading to a Na/Cl– SO4 brine (but starting in the Ca>SO4 area) and the other (line B) ending in a Na – Ca / Cl brine. Owing to the large number of inflows (354), it is not possible to show the theoretical evolutionary path of each of them and to compare it to the actual composition of the lake or pond it feeds. Actual evolutionary paths after the gypsum divide are in
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Fig. 16. SO4 and SO4 + Alk/2 concentration vs. Ca for inflow waters and lake brines in Chilean salars. Inset flow diagrams show the evaporative concentration paths followed by each group of waters. The evolutionary path of each inflow water has been determined with the EQL/EVP evaporation program. In contrast, lakes are simply differentiated according to their Ca/SO4 ratio. Many waters following the SO4-rich path are actually in the Ca>SO4 field. The early precipitation of calcite lowers the Ca concentration and reverses the Ca/SO4 ratio. Line A shows such an evolutionary path. Line B shows a typical Ca-rich path.
general agreement with the calculated paths. We could not detect reversion from one path to the other, as observed after the Mg salts divide when waters reverse from the alkaline path to the sulfatealkaline path due to eolian sulfur. The gypsum divide
occurs at rather high Ca and SO4 concentrations once their ratio can no longer be easily modified. Nevertheless, the mixing of brackish waters of different Ca/SO4 ratio may produce unpredictable end brines.
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In total, 243 inflow waters follow either the sulfate-alkaline or the sulfate-neutral path which ends up as SO4-rich solutions in 143 lakes and ponds, while 60 inflows follow the calcic path leading to Ca-rich solutions feeding 82 lakes and ponds (Fig. 16). In addition to the 243 inflows following the sulfate paths, 52 inflows that theoretically should follow the alkaline path up to carbonate brines also end up as sulfate-rich solutions after reverting to the sulfate-alkaline path (Fig. 15). All the 225 lakes and ponds (143 + 82) are distributed in 52 salars. Table 1 shows to which chemical group each salar belongs. Approximately two-thirds of the salars (34 out of 52, SO4 in Table 1) contain only brackish and saline waters high in SO4 and low in Ca. Twelve salars belong to the Ca-enriched group (Ca) and seven contain both Ca-rich and SO4-rich solutions. Several salars of the Ca-rich group have a few SO4-rich brackish ponds. However, their most concentrated brines are Ca-rich. The symmetric situation is not observed in SO4-rich salars, where Ca-rich brines are completely lacking.
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Salar de Atacama, which is the largest in Chile, contains both Ca-rich and SO4-rich brines within the halite nucleus (Risacher and Alonso, 1996). The western part of the nucleus, close to the Cordillera de Domeyko where sedimentary rocks are predominant, is filled with Ca-rich brines, while the eastern part close to volcanic formations of the Cordillera is filled with SO4-rich brines. The Ca vs. SO4 diagram shows very clearly the gypsum divide and the relation between lithology and brine composition (Fig. 17; data from Moraga et al., 1974; Ide, 1978). As can be seen in the upper diagram of Fig. 16, it is not possible to use the initial Ca/SO4 ratio of the inflow water to predict which evaporative path will be followed after the gypsum divide. In contrast, as shown in the lower diagram, Ca vs. SO4 + Alk/2, the isoconcentration line Ca = SO4 + Alk/2 separates perfectly the two groups of inflow waters. Those with initial Ca>SO4 + Alk/2 follow the Ca-rich path, and those with SO4 + Alk/2>Ca follow the SO4-rich path. This is obvious if Ca>SO4 + Alk/2: whatever the mass of CaCO3 and CaSO4 precipitated, Ca will remain in
Fig. 17. SO4 vs. Ca diagram illustrating the gypsum divide in salar de Atacama brines. (These data are not included in Fig. 16.) Points are distributed along a tilted T. The ascending branch (+) corresponds to waters undersaturated with respect to gypsum. Waters saturated with gypsum plot along the approximate gypsum saturation line and are divided into two groups: SO4-rich brines (circles) above the isoconcentration line Ca = SO4, and Ca-rich brines (squares) below. The geograpical distribution of the three groups of solutions is shown in the Atacama outline map. Analyses are from Moraga et al. (1974) for the SO4 – Ca diagram and from Ide (1978) for brine distribution in the salt nucleus. Note that the approximate gypsum saturation line has no rigorous chemical meaning, inasmuch as the true saturation line can only be drawn in an activity diagram.
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Table 5 Theoretical evaporative paths and final brine composition predicted from the initial composition of inflow waters Concentrations in inflow water
Evaporative path
Brine group
Alk>2Ca and Alk>2Ca + 2Mg Alk>2Ca and Alk < 2Ca + 2Mg Alk < 2Ca and SO4 + Alk/2>Ca Alk < 2Ca and SO4 + Alk/2 < Ca
alkaline (I – IA)
Na/CO3 – Cl
sulfate-alkaline (I – IIA – III) sulfate-neutral (II – III) calcic (II – IV)
Na/SO4 – Cl Na/SO4 – Cl Na – Ca/Cl
Path numbers as in Fig. 14. Concentrations are in mol/l or mmol/l.
excess. However, this rule cannot be rigorously demonstrated in the case SO4 + Alk/2>Ca. If the precipitation of calcite is not occurring and if Ca>SO4, then the Ca-rich path could be followed. The most common way to prevent, or to delay, the precipitation of calcite is to increase the content of CO2 gas in the
water. This happens in confined aquifers where the water is not in contact with the atmosphere and therefore does not evaporate. In evaporitic environments, the CO2 dissolved in waters is close to equilibrium with that of the atmosphere. Thus, the ratio (SO4 + Alk/2)/Ca of the initial water may be used with good confidence to predict which of the Ca-rich or SO4-rich path will be followed by an evaporating water after the gypsum divide. Table 5 summarizes the relations between the initial composition of inflow waters and their evaporative path. 9.3.3. Relation between evolutionary path and temperature We have plotted in Fig. 18 the ratio (SO4 + Alk/2)/ Ca of all springs and underground waters against the difference between water temperature and mean annual air temperature (seeps and rivers are not considered). Positive Y-axis values (log scale) correspond to the SO4-rich path and negative values
Fig. 18. Diagram showing the relation between temperature and the two main evolutionary paths SO4-rich and Ca-rich for springs and underground waters. Waters are differentiated according to their salt content. Most hot inflow waters follow the SO4-rich path. Note that the region labeled the ‘‘SO4-rich path’’ include the waters that have reversed their theoretical evolutionary path from alkaline-rich to sulfate-rich after probable acidification by sulfur oxidation.
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correspond to the Ca-rich path. Almost all high thermal inflow waters (Twater Tatm>20 jC) belong to the SO4-rich group. Only 3 out of 22 hot inflow waters follow the Ca-rich path. It must be emphasized that the evolutionary path is the only parameter for which some relation with temperature has been detected so far. None of the dissolved components presents any relation with temperature (see Section 7.7). 9.3.4. Discrepancy between lithology and brine evolution The geographical distribution of SO4-rich and Carich salars is shown in Fig. 19. According to the previous discussion, salars located in volcanic basins are expected to contain SO4-rich brines, while those surrounded by sedimentary rocks should be filled with Ca-rich brines, as perfectly illustrated in salar de Atacama [13]. In addition to Atacama, two other salars are located in the Pre-Andean Depression. where sedimentary formations of the Precordillera predominate: Imilac [29] and Punta Negra [30]. Both belong to the Ca-rich type. All other salars, except Pintados [5], are in volcanic basins of the Western Cordillera and are expected to be of the SO4-rich type as a result of the abundance of native sulfur in most of the volcanoes. This is effectively observed in the north of the study area (18 – 21jS latitude). Salar de Pintados [5], in the Central Valley, is mainly fed by inflows originating in the volcanic Western Cordillera. In the central area (21 –24jS latitude), both types are intermingled. No simple geological argument can explain the difference from one basin to the other. In the south of the study area (25 – 27jS latitude), salars are predominently of the Ca-rich type close to the Precordillera and of the SO4-rich type eastward in the volcanic Cordillera. The chemistry of Ca-rich salars located in the volcanic Western Cordillera seems to be controlled by sedimentary formations. If some of them are adjacent to the Precordillera (Pedernales [44], Maricunga [52]), others are well inside the Western Fig. 19. Outline map of north Chilean Andes showing the chemistry of each salar: SO4-rich, Ca-rich or both. In the south, Ca-rich salars predominate close to the Precordillera, while SO4-rich salars are found eastward within the Cordillera (compare with Fig. 12). Note the slightly exaggerated horizontal scale.
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Cordillera, up to 50 km from its western boundary (Aguas Calientes 2 [21]; Wheelwright [49]). It is difficult to conceive a major contribution of sedimentary rocks, either from the west or from depth, in salars located in such massive volcanic formations. A much better explanation is the recycling of present infiltrating brines from the salars themselves (see Section 7). However, such an explanation only pushes the problem further back in time. When did the waters acquire their Ca-rich composition? The drainage basins of the Ca-rich salars, now mostly or exclusively volcanic, must have been sedimentary in the past, before the establishment of recent Pleistocene volcanoes and ignimbrite sheets. Consequently, these salars should be notably old, of Pleistocene age or even earlier. Stoertz and Ericksen (1974) have already
observed that many salars in the Western Cordillera have been tilted downward to the northwest during the Quaternary. Salares de Pedernales [44] and Maricunga [52] are precisely among the most asymmetric ones, which points to an older age than that of nontilted salars within the Western Cordillera. The ancient Carich underground waters may have been continuously recycled up to now. The alteration of volcanic rocks progressively shifts the chemical type of the solutions. Both Pedernales and Maricunga belong to the mixed Ca/SO4 salar group. Inflow waters result from the mixing in variable proportion of recycled brines with dilute weathering solutions. The weathering of sulfur-rich volcanic rocks increases the SO4 content, and then the SO4/Ca ratio, of inflow waters. Alternatively, if the
Fig. 20. Very schematic representation of the recycling and mixing processes responsible for the chemical shift from Ca-rich to SO4-rich brines in Chilean salars. For the sake of simplicity, we have restricted the cycle to one basin. Actually, deep water circulation may transport infiltrating salts in several distant basins. Size of symbols only reflects the proportion of Ca and SO4 in each solution. The salar brine is of the Ca-rich type, which means that the concentration of Ca is markedly higher than that of sulfate (as suggested by the thick bold characters used for Ca). The infiltrated brine has the same composition. Heat flow recycles this Ca-rich brine along the Ghyben – Herzberg mixing zone, in contact with fresh waters that have leached volcanic rocks of the drainage basin. Sulfur-rich volcanoes produce SO4-rich waters which mix with Ca-rich recycled brines. Thus, the Ca/SO4 ratio of the mixed solution is lower than that of the salar brine (spring 1). The gap between Ca and SO4 concentration is reduced. The weathering of sulfur-poor volcanoes produces alkaline waters. Calcite is a common weathering product whose precipitation lowers the Ca concentration of the water. Here again, the mixing of Ca-poor waters with Ca-rich infiltrated brines reduces the gap between Ca and SO4 concentration in the resulting inflow water (spring 2). Therefore, on the one hand, inflows tend to reduce the excess of Ca over SO4 in the salar brine; on the other hand, the evaporative concentration amplifies the difference between Ca and SO4 in the brine. A long time is likely to be necessary to shift the Ca-rich brine composition to a SO4-rich composition.
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sulfur content of volcanic rocks is low, then the weathering solutions are enriched in HCO3 (see Section 9.1), which also reduces the Ca concentration of the mixed inflow waters due to calcite precipitation. Here again, the SO4/Ca ratio of the inflow water increases. However, the evaporative concentration in the basin amplifies the difference on behalf of the most concentrated component (Ca). Thus, the SO4/Ca ratio of the infiltrating brine is much lower (more Ca) than that of the input water. This amplification mechanism may delay for a long time the shift from the Carich to the SO4-rich type. Eventually, SO4 overtakes Ca in spring waters and the amplification process then favors the rapid increase of SO4 in lake brines. Fig. 20 summarizes the whole cycle.
10. Concluding remarks Three main sources of salts have been detected in inflow waters of Chilean salars. Atmospheric inputs, through precipitation and dry fallout, affect the whole area but are noticeable only in a few salars fed by very dilute inflows [below 100 mg/l total dissolved solids (TDS)]. Most of the minor components in very dilute inflows, and a significant part of the major ones, stem from sea salts and desert dust, especially NO3, Br and As. Volcanic rock alteration is the main source of salts in inflow waters with concentration ranging roughly from 100 to 600 mg/l TDS. However, the average salt content of all inflows (3230 mg/l) is much higher than that derived from the mere alteration of volcanic rocks. Brine recycling is the main source of salts in most salars. Saline lakes are under steady-state regime. The residence time of conservative components ranges from a few years to some thousands years. Most of the dissolved salts brought in the lakes by inflow waters are lost by infiltration through bottom sediments and return to the hydrologic cycle. The infiltrating brines provide the major part of dissolved salts to underground waters. The dissolution of surficial salt crusts also contributes to the salt load of deep waters. Owing to the complexity of water circulation in fractured volcanic formations, infiltrating brines may be transported to adjacent basins, or even further. Thermal convection due to the high heat flow is the likely driving force that moves the waters. Thermal heating affects about 78% of springs and
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underground waters. Thermal heating may have different results below the salar and below the mountains. Below the salar, heating could tend to oppose the downward brine flux and favor recycling in the same salar. In contrast, below the volcanic formations, the heating could increase the loss of water from the basin to other basins (see Fig. 20). Ancient evaporites, especially halite and gypsum, are present in depth below the volcanic formations. However, the Cl/Br ratio of halite is much higher than that of inflow waters, which precludes the dissolution of deep salt rock in underground waters. On the contrary, the Cl/ Br ratio confirms the mixing of present lake brines with snow meltwaters as the origin of saline inflows. Only ancient gypsum seems to contribute some Ca and SO4 to inflow waters. Stable isotopes also supports the mixing hypothesis: the d2H and d18O values of inflow waters plot along a straight line, probably a mixing line, joining snow meltwaters to lake brines. Chilean salars in the Andean Cordillera cannot be strictly considered as closed basins or terminal lakes. Most of the salts entering the salars are lost by infiltration. Only poorly soluble salts, such as calcite, Mg salts and gypsum, remove dissolved components which slowly accumulate in saline sediments. The high concentration of many inflows does not reflect an intense material transfer by weathering, evaporite dissolution or hydrothermal alteration. Nor does it mean a high accumulation rate of salts in the basins. Actually, true material transfer is mostly restricted to atmospheric inputs and volcanic rock weathering. This is intuitively understandable: if all inflowing salts would accumulate in Andean salars, very thick salt crusts should be observed. Thick modern salt crusts are not found in this region. According to the model of brine evolution observed in most saline lakes worldwide, three main groups of brines may be generated by evaporation and chemical divide (Hardie and Eugster, 1970): alkaline brines (Na/HCO3 – CO3 – Cl, pH>9, traces of Ca and Mg), sulfate-rich brines (Na/SO4 –Cl, low Ca, pH < 9) and Ca-rich brines (Na – Ca/Cl, low SO4, pH < 9). In northern Chile, alkaline salars are almost completely lacking. Only one brackish pond out of 227 was found to be alkaline. All others belong to the SO4-rich and Ca-rich groups. This paucity is due to the abundance of native sulfur in the Western Cordillera and to atmospheric deposition of gypsum-rich desert dust
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from the Central Valley. The oxidation of sulfur produces SO4 and acidifies the inflow waters, reducing drastically their carbonate content. Gypsum dust enriches the waters in Ca, which conversely lowers their carbonate contents because of the very low solubility of calcite. Therefore, the (SO4 + Alk/2)/Ca ratio was found to be a valuable parameter to class and investigate Chilean salars. The abundance of native sulfur in volcanoes of the Western Cordillera leads to sulfate-rich inflow waters, which in turn produce sulfate-rich brines by evaporative concentration. Conversely, sedimentary rocks of the Precordillera contain Ca minerals (calcite, dolomite, gypsum) whose weathering leads preferentially to Ca-rich waters and Ca-rich brines. Small differences of SO4 and Ca concentration in dilute inflows are amplified in the final brines by evaporative concentration. This overall relation between lithology and brine composition is remarkably illustrated in the largest Chilean salar: the salar de Atacama. However, several salars in volcanic basins of the Western Cordillera belong to the Ca-rich group. Their chemistry may have been inherited from the weathering of ancient sedimentary formations during the late Tertiary or the Pleistocene. Later on, ignimbrite sheets and lava flows may have covered the sedimentary formations. Such a process may well account for Carich salars adjacent to the Precordillera, at the western fringe of the volcanic cordillera. Most of the salt load of inflow waters does not stem directly from rock alteration in the drainage basins, but from lake brine recycling. The recycling process is likely to be as old as the salars themselves. Infiltration rates tend to slowly decrease with time because of the progressive choking of bottom sediments by clay and silt particles. Thus, brine leakage was probably more effective in past times. Nonequilibrium conditions seem to control important features of Chilean salars. Old salt crusts deposited under very arid climate are presently undergoing dissolution in the semiarid Cordillera. The Ca-rich composition of many waters and brines does not reflect the volcanic lithology of the Western Cordillera. The global hydrogeochemical cycle is not yet equilibrated with the presently dominant volcanic lithology. The quantity of weathering material entering the hydrological cycle is much lower than the quantity of recycled salts.
A complex correlation between lithology, thermalism and salar chemistry is hinted at in the south of the study area. Ca-rich salars are preferentially located close to the sedimentary Precordillera and their inflow waters undergo little thermal influence. Conversely, SO4-rich salars tend to be located well inside the volcanic Western Cordillera and are largely fed by thermal waters. This is consistent with the sedimentary to volcanic transition from west to east associated to the increase of heat flow in the same direction. Obviously, many factors obscure this overall correlation: the sedimentary formations contain volcanic and volcano-sedimentary units. The boundary between Precordillera and Western Cordillera is not so precisely delimited as shown in our schematic maps. Thermal waters may be cooled by mixing with fresh meteoric waters. Deep circulation in fractured eruptive rocks can be very complex. The northern and central areas seem to be particularly affected by these drawbacks. Dilute meteoric waters are more contaminated by salt recycling in areas containing many salars. Salts lost from one salar may be widely redistributed in several other basins. As a result, exploration for fresh water should be focused in areas devoid of salars. In salar basins, the water table is close to the topographic surface, which reduces significantly the drilling costs. The drawback is that such waters are often of poor quality. Outside salar areas, waters are much deeper, but probably less affected by salt contamination.
Acknowledgements We want to express our deep gratitude to the institutions and companies that gave us invaluable and unselfish logistic assistance: the regional Bureaus of the Direccio´n General de Aguas (Regiones I, II, III); Eje´rcito de Chile; Minera Escondida; Codelco Chile: Division Chuquicamata and Division Salvador; Minera Mantos Blancos; Corporacion Nacional Forestal; Minera Quiborax; Minera Don˜a Ine´s de Collahuasi; and Minera Zaldivar. Thanks are due to Professor Abraham Lerman and to an anonymous reviewer who made constructive reviews of the manuscript. This is EOST contribution no. 2003.403-UMR7517.
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Appendix A (continued)
Appendix A Cl (mmol/l) and Br (Amol/l) of waters and brines of Chilean salars. Other components in Risacher et al. (1999). N = snow, S = spring, P = seep, R = river, L = lake, U = underground water.
Name
Type
CL/BR
NEV-1 NEV-2 NEV-3 NEV-4 NEV-5 NEV-6 NEV-8 NEV-9 AC1-1 AC1-4 AC1-6 AC1-8 AC1-11 AC2-1 AC2-4 AC2-7 AC2-10 AC2-13 AC2-15 AC3-1 AC3-4 AC3-7 AC3-10 AC3-12 AC3-13 AC3-14 AC3-15 AC4-1 AC4-2 AC4-3 AC4-4 AC4-5 AC4-6 AC4-7 AC4-8 AC4-9 AC4-10 AGI-1 AGI-1A AGI-2 ALC-1 ALC-2 ALC-3 ALC-4 ALC-5 ALC-6
N N N N N N N N S S L L S S P S P P L P L S P P P S L P L L P L L L P S S L L L U S S S S L
0.015/0.05 0.021/0.09 0.018/0.13 0.029/0.10 0.025/0.08 0.003/0.02 0.003/0.04 0.005/0.02 408/56 102/18 1820/280 2020/210 17.3/7.3 19.4/2.6 35.5/7.7 40.0/12 217/28 35.2/13 165/45 75.0/14 361/46 28.0/11 126/13 22.3/4.2 116/40 34.1/14 48.0/18 9.39/5.0 380/190 88.3/43 27.1/12 271/130 892/460 4930/2400 25.2/16 5.22/2.8 16.9/10 2960/230 3200/230 5880/150 2.52/2.7 0.169/0.49 0.168/0.57 0.197/0.30 0.184/0.40 1030/370
Name
Type
CL/BR
ALC-7 ALC-8 ALC-9 ALC-10 AMA-1 AMA-2 AMA-3 AMA-4 AMA-5 AMA-6 ASC-1 ASC-2 ASC-3 ASC-4 ASC-7 ASC-9 ASC-10 ASC-13 ASC-14 ASC-15 ASC-16 ASC-17 ASC-18 ASC-19 ASC-20 ASC-21 ASC-22 ASC-23 ASC-24 ASC-25 ASC-26 ASC-27 ASC-28 ASC-29 ASC-30 ASC-31 ASC-32 ASC-33 ASC-34 ASC-36 ASC-37 ASC-38 ASC-39 ASC-40 ASC-41 ASC-42 ASC-43 ASC-44 ASC-45 ASC-46 ASC-47 ATA-1 ATA-2
P L L S L L P L S L U U U U U U U S S U U U L S L L U L U L S L S L S L L S L S S S L L L P L S L L U R S
2.03/2.7 14.9/24 58.6/110 0.328/1.0 538/31 284/17 129/9.2 3410/200 279/52 1420/68 32.7/18 8.88/1.5 25.3/15 30.5/9.4 109/8.6 35.5/9.0 60.1/10 0.087/0.03 0.081/0.05 25.0/14 3.01/1.9 47.2/13 211/43 122/7.9 124/9.0 287/21 31.6/7.7 91.3/19 31.4/8.7 67.5/21 30.9/6.0 64.7/15 56.3/8.7 151/23 54.4/10 136/21 657/150 55.8/9.7 199/41 49.1/18 33.9/17 27.8/13 266/120 107/24 1650/550 28.4/1.7 916/170 3.69/0.75 583/120 245/68 30.3/10 11.8/0.48 2.32/0.28 (continued on next page)
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Appendix A (continued)
Appendix A (continued)
Name
Type
CL/BR
Name
Type
CL/BR
ATA-3 ATA-4 ATA-5 ATA-6 ATA-7 ATA-8 ATA-9 ATA-10 ATA-11 ATA-12 ATA-13 ATA-14 ATA-15 ATA-16 ATA-17 ATA-18 ATA-19 ATA-20 ATA-21 ATA-22 ATA-23 ATA-24 AZU-1 AZU-2 AZU-3 AZU-4 AZU-5 AZU-6 AZU-7 BAY-1 BAY-2 BRA-1 BRA-2 BRA-3 BRA-4 CAR-1 CAR-2 CAR-3 CAR-4 CAR-5 CAR-6 CAR-7 CAR-8 CAR-9 CAR-10 CAR-11 CAR-12 CAR-13 CAR-15 CAR-16 CAR-17 CAR-18 CAR-19 CAR-20
R R U R R R R R R R R R R S R R P P P U S S P L L L L L U L P P L P L S S L L S L L L L L L S L S L L S L L
7.10/1.6 17.0/1.3 27.0/18 1.50/1.0 8.50/5.4 2.03/5.0 1.25/1.0 6.50/3.8 2.50/1.6 25.0/11 43.0/11 15.4/10 15.0/10 15.0/10 3.50/2.1 0.480/0.42 10.5/7.4 304/40 25.3/11 52.9/14 7.76/4.2 0.041/0.08 27.4/6.4 893/78 2190/280 120/23 292/56 5700/870 4.41/1.1 28.9/23 15.9/7.1 35.9/15 1810/790 21.8/11 47.7/24 0.269/0.37 0.272/0.32 1030/59 3020/130 47.1/3.1 2020/95 5700/530 2900/190 31.7/2.3 190/8.8 1250/70 126/13 1800/190 17.4/2.3 5490/130 463/87 11.4/2.6 273/48 36.8/7.6
CAR-21 CAR-22 CAR-24 CAR-25 CAR-26 CHR-1 CHR-2 CHR-3 CHR-4 CHR-5 CHR-6 CHR-7 CHR-8 COP-1 COP-2 COP-3 COP-4 COP-5 COP-6 COP-7 COP-8 COP-9 COP-10 COP-11 COP-13 COP-14 COP-15 COP-16 COP-17 COP-18 COP-19 COP-20 COP-21 COP-22 COP-23 COP-24 COT-1 COT-2 COT-3 COT-4 CPR-1 CPR-2 CPR-3 CPR-4 CPR-5 CPR-6 CPR-7 ESC-1 ESC-2 ESC-3 ESC-4 ESC-5 ESC-6 FRA-1
L S P S L L R S S L S L S S S U P S L U U U L L L L L U S L L L L U U U S L L L L S L S L L L L L P P L L R
90.5/11 6.74/2.2 109/8.4 419/22 1250/69 2.71/0.57 0.091/0.09 0.007/0.05 0.045/0.06 0.087/0.08 0.046/0.05 1.74/0.23 0.016/0.02 1.72/1.8 43.9/25 0.269/0.38 5.30/2.0 4.89/2.2 461/180 28.5/4.3 0.777/0.78 21.2/2.6 461/160 1430/29 32.3/12 121/42 5.60/2.4 20.9/2.2 13.7/3.6 807/190 187/46 43.6/10 4410/830 3.01/1.6 1.13/1.0 0.739/0.89 0.168/0.14 0.838/0.32 0.903/0.16 2.55/1.3 197/25 106/17 334/40 102/16 847/76 1190/100 3770/280 41.1/33 199/180 18.7/14 30.5/30 347/330 71.6/59 0.483/0.17
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293 Appendix A (continued)
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Appendix A (continued)
Name
Type
CL/BR
Name
Type
CL/BR
FRA-2 FRA-3 FRA-4 FRA-5 FRA-6 FRA-7 FRA-8 FRA-9 FRA-10 FRA-12 FRA-13 FRA-14 FRA-15 FRA-16 FRA-17 FRA-18 GOR-1 GOR-2 GOR-3 GOR-4 GOR-5 GOR-6 GOR-7 GOR-8 GOR-9 GOR-10 GOR-11 GOR-12 GRA-1 GRA-2 GRA-3 GRA-4 HCO-1 HCO-2 HCO-3 HCO-4 HCO-5 HCO-6 HCO-7 HCO-8 HCO-9 HCO-10 HCO-11 HCO-12 HCO-13 HCO-14 HCO-15 HCO-16 HCO-17 HCO-18 HEL-1 HEL-3 HEL-5 HEL-6
P R L P S L S L S P S S L R R S R P L L L P P L L P L L P L L L S R R R R U S R U R S S L L L L L L S L S S
2.73/1.3 0.023/0.03 4810/2700 0.608/0.41 0.405/0.07 429/170 0.630/0.10 97.0/37 3.41/2.0 11.4/7.5 1.02/0.49 0.339/0.30 20.1/3.2 3.05/2.5 0.322/0.14 83.1/88 3.36/0.72 54.5/41 111/86 258/48 1350/230 5.63/2.0 112/8.6 104/100 506/270 46.4/52 3450/110 3490/80 119/12 305/30 2120/170 440/47 0.819/1.2 0.145/0.25 0.269/0.64 0.256/0.47 0.748/0.72 2.30/2.6 0.100/0.40 0.088/0.36 0.576/0.73 0.390/0.64 1.43/0.98 0.916/1.3 140/78 222/140 80.2/47 12.5/8.0 246/150 909/470 1.66/0.24 5050/940 0.250/0.31 0.109/0.18
HEL-7 HEL-8 HEL-9 IGN-1 IGN-3 IGN-4 IGN-5 IMI-1 IMI-2 IMI-3 IMI-5 IMI-6 IMI-7 IMI-8 IMI-8D INF-1 INF-2 INF-3 INF-4 INF-5 INF-6 ISL-1 ISL-2 ISL-3 ISL-4 ISL-5 ISL-6 ISL-7 ISL-8 ISL-9 ISL-10 ISL-11 ISL-12 ISL-13 ISL-14 ISL-15 ISL-16 ISL-17 ISL-18 ISL-19 JIL-1 JIL-2 JIL-3 JIL-4 JIL-5 JIL-6 LAA-1 LAA-2 LAA-3 LAA-4 LAA-5 LAA-6 LAA-7
S L L P L L L U L U L L L P U P L L L L P S L L P L L L U S L L L S L L S L S L L P L L L P U R R R S R S
0.537/0.35 23.7/5.6 14.9/22 4.00/1.1 17.0/1.8 12.5/1.9 39.6/5.1 2.70/2.0 3.24/2.1 55.5/7.8 1410/8.4 713/8.9 507/11 150/5.7 181/6.9 15.3/6.2 142/41 5540/250 1230/97 280/28 119/19 276/2.2 5310/28 1410/8.0 102/1.1 1540/11 425/2.2 818/11 155/3.6 262/4.3 658/11 2130/21 5070/22 494/2.5 2920/16 1220/8.0 226/2.5 5030/38 88.3/1.5 2470/52 177/100 9.93/7.3 115/67 123/67 215/110 32.3/13 0.101/0.29 1.14/1.1 0.964/0.85 1.52/1.4 0.682/0.45 0.956/0.51 0.246/0.32 (continued on next page)
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Appendix A (continued)
Appendix A (continued)
Name
Type
CL/BR
Name
Type
CL/BR
LAA-8 LAC-1 LAC-4 LAC-5 LAC-7 LAV-1 LAV-2 LAV-3 LAV-4 LAV-5 LAV-6 LAV-7 LAV-8 LAZ-1 LAZ-2 LAZ-3 LAZ-4 LEJ-1 LEJ-3 LGN-1 LGN-2 LGN-3 LGN-4 LGN-5 LGU-1 LGU-2 LGU-3 LOY-1 LOY-5 LOY-9 LOY-13 LOY-16 LOY-20 LOY-24 LOY-28 LOY-32 LOY-35 LOY-39 LOY-40 LOY-41 MAR-1 MAR-2 MAR-3 MAR-4 MAR-5 MAR-6 MAR-7 MAR-8 MAR-9 MAR-10 MAR-11 MAR-12 MAR-13 MAR-14
R P U L P L R P L L S S R S L L P L P P L P L S S L U R P P R P L P P U P L S S S R R R R U R R R L L P P L
0.268/0.38 5.43/1.8 5.22/1.5 325/45 24.2/5.9 2580/1300 7.60/3.8 78.7/72 56.4/58 520/220 41.1/16 8.02/3.1 24.0/12 37.5/28 1580/990 166/110 28.0/15 274/130 3.14/1.0 22.3/9.3 73.9/41 100/54 232/160 28.6/30 0.906/1.2 4.53/3.7 0.268/0.88 38.4/3.5 230/19 1.50/0.53 0.515/0.33 3.38/0.68 64.3/6.2 108/20 59.0/4.8 57.2/10 62.0/5.5 4170/65 1.11/0.51 0.489/0.26 0.501/0.22 0.856/0.63 5.29/3.1 29.5/17 1.00/0.07 2.44/2.4 12.2/7.3 0.096/0.04 2.20/1.2 61.0/14 3.48/2.0 35.3/19 19.1/9.7 251/32
MAR-15 MAR-16 MAR-17 MAR-18 MAR-20 MAR-21 MAR-22 MAR-23 MAR-24 MAR-25 MAR-26 MAR-27 MAR-28 MAR-29 MAR-30 MIC-1 MIC-2 MIC-3 MIC-4 MIC-5 MIC-6 MIC-7 MIC-8 MIC-10 MIC-13 MIC-14 MIC-15 MIN-1 MIN-3 MIN-6 MIN-8 MIS-1 MIS-3 MIS-5 MIS-8 MIS-9 MUE-1 MUE-4 PAJ-1 PAJ-2 PAJ-3 PAJ-4 PAJ-5 PAJ-6 PAJ-7 PAJ-8 PAJ-9 PAJ-10 PAJ-11 PAR-1 PAR-2 PAR-3 PAR-4 PAR-5
P U U U L P L L L L P S U P R S R S L L L L U U U S L L L P L P L P L U P L S L U L L L P S L L L P L L P L
13.8/2.3 62.1/25 26.1/15 0.838/0.54 130/13 38.5/6.3 66.8/12 159/38 1420/280 5750/1300 22.7/3.9 1.12/0.85 1.20/0.27 0.185/0.05 10.6/4.0 0.138/0.23 0.594/0.48 0.664/0.60 3.95/2.4 181/98 20.9/8.8 50.0/22 0.511/0.57 0.661/0.75 0.397/0.65 0.157/0.44 5.21/5.5 74.6/30 67.1/24 23.4/9.0 62.6/24 1.57/0.90 32.4/13 3.86/1.4 34.0/12 0.089/0.37 364/38 1090/130 240/38 907/140 371/53 723/120 4250/620 2820/270 381/35 193/59 1890/670 3720/820 2390/690 139/3.6 568/15 3710/87 254/6.8 5290/180
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293 Appendix A (continued)
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Appendix A (continued)
Name
Type
CL/BR
Name
Type
CL/BR
PAR-6 PAR-7 PED-1 PED-2 PED-3 PED-4 PED-5 PED-6 PED-7 PED-9 PED-10 PED-11 PED-12 PED-13 PED-14 PED-15 PED-16 PED-17 PED-18 PED-19 PED-20 PED-21 PED-22A PED-22 PED-23 PED-25 PED-26 PED-27 PED-28 PED-29 PIE-1 PIE-2 PIE-3 PIE-4 PIE-5 PIE-6 PIE-7 PIE-8 PIE-9 PIE-10 PIE-11 PIE-12 PIE-13 PIE-14 PIN-1 PIN-2 PIN-3 PIN-4 PIN-5 PIN-6 PIN-8 PIN-9 PIN-10 PIN-11
S L S R P R S S U U P L P S L L P S S L U U L L L U U U U U P L L P L L L P L L S L P L U U U S U U S R S U
133/4.3 708/23 28.2/22 29.9/13 2.87/1.1 8.62/2.5 6.43/0.47 61.0/6.1 1.81/0.93 0.070/0.05 2.89/1.6 3640/190 1310/68 40.3/2.4 1790/50 5450/97 245/8.7 4.87/5.6 0.799/0.67 5410/180 0.157/0.16 0.148/0.14 175/35 510/97 1010/120 0.252/0.27 35.2/12 33.3/12 40.9/16 35.5/13 41.6/8.2 353/75 97.5/22 84.4/29 304/110 492/130 269/49 38.8/6.9 1220/190 163/38 30.0/6.9 419/170 38.8/7.3 141/26 2.10/2.3 3.70/3.1 2.60/2.7 6.00/5.0 5.96/5.3 17.4/17 5.26/4.4 74.5/45 1.80/2.1 23.4/18
PIN-12 PIN-13 PIN-14 PIN-15 PIN-16 PIN-17 PIN-18 PIN-19 PIN-20 PIN-21 PIN-22 PIN-23 PIN-24 PIN-25 PIN-26 PIN-27 PIN-28 PIN-29 PIN-30 PIN-31 PIN-32 PIN-33 PIN-34 PIN-35 PIN-36 PIN-37 PIN-38 PIN-39 PIN-40 PIN-41 PIN-42 PIN-43 PIN-44 PIN-45 PSA-1 PSA-5 PSA-8 PSA-11 PUN-1 PUN-2 PUN-3 PUN-4 PUN-5 PUN-6 PUN-7 PUN-8 PUN-9 PUN-10 PUN-11 PUN-12 PUN-13 PUN-14 PUN-15
S U S U U U S S S S S S S U U U U U R R S S R U L R S S U L U L L L R L P P L L L P L L L U L P U S S S U
0.740/1.4 0.931/1.6 2.49/2.4 3.12/3.6 19.4/16 4.09/4.5 1.22/2.3 1.47/2.7 1.18/0.44 1.15/2.2 3.01/4.5 1.14/2.2 3.98/6.5 6.82/6.7 19.1/17 21.9/19 23.6/21 6.97/6.5 32.3/29 1.06/2.1 1.88/2.9 1.51/2.0 7.69/5.8 10.0/7.9 935/120 3.65/3.4 0.391/1.8 0.406/1.7 7.08/6.6 236/39 12.3/9.3 4390/680 1440/190 115/24 0.983/0.61 343/230 6.36/3.5 0.559/0.20 309/92 195/57 162/47 170/11 294/21 4010/120 155/3.8 112/2.2 706/17 238/9.8 33.6/11 2.12/2.4 5.17/2.8 35.6/31 37.9/32 (continued on next page)
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Appendix A (continued)
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Name
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CL/BR
PUN-16 PUN-17 PUN-19 PUN-21 PUN-22 PUN-23 SUR-1 SUR-2 SUR-3 SUR-4 SUR-5 SUR-7 SUR-8 SUR-9 SUR-10 SUR-11 SUR-12 SUR-13 SUR-14 SUR-15 SUR-16 SUR-17 SUR-18 SUR-19 SUR-20 SUR-21 SUR-22 SUR-23 SUR-24 SUR-25 SUR-26 TAR-1 TAR-3 TAR-5 TAR-6 TAR-9 TAR-11 TAR-12 TAR-13 TRI-1 TUY-1 TUY-4 WHE-1 WHE-2 WHE-3 WHE-4 WHE-5 WHE-6 WHE-7 WHE-8 WHE-9
U R U U S R R S P S P S L S P P S L S P S L S L S L L L L L S S L R P L S S R L R L R L S L L L P L P
43.9/36 43.4/32 4.84/0.77 6.48/4.8 0.184/0.28 0.491/0.35 0.141/0.19 47.9/36 3.27/2.9 1.49/1.2 6.52/4.0 0.653/0.97 19.0/14 39.3/32 23.2/19 0.793/0.79 1.20/1.0 1040/680 0.514/0.85 0.526/0.57 6.12/4.6 614/380 0.375/0.53 459/300 226/170 967/680 50.6/33 4380/2700 1970/1300 89.3/66 0.286/0.59 8.60/4.0 34.8/15 1.60/1.1 7.25/3.5 2760/1100 9.18/5.0 26.4/13 15.7/6.8 131/14 88.9/9.7 2120/160 100/11 2750/270 48.7/5.2 329/42 2480/290 708/96 96.1/12 175/25 20.2/3.2
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aquifer, Rainer Mesa, Nevada. U.S. Geological Survey WaterSupply Paper 1535-Q. Yechieli, Y., Wood, W.W., 2002. Hydrogeologic processes in saline systems: playas, sabkhas, and saline lakes. Earth-Science Reviews 58, 343 – 365. Yonts, W.L., Giese, G.L., Hubbard, E.F., 1973. Evaporation from Lake Michie, North Carolina, 1961 – 71. U.S. Geological Survey Water-Resources Investigations 38-73, 27 pp. Francß ois Risacher is a graduate of the Ecole Nationale Supe´rieure de Ge´ologie de Nancy (1971). He holds a PhD in Geochemistry from the University Louis Pasteur of Strasbourg (1992). He is a Senior Researcher at the Institut de Recherche pour le De´ veloppement (a French cooperation agency formerly known as Orstom). For more than 20 years, he carried out applied and fundamental research on evaporitic basins in Bolivia and Chile, focusing his interest on economical resources and geochemistry of saline lakes and salt crusts. Hugo Alonso is a Chemical Engineer from the Universidad Cato´lica del Norte (UCN), Chile (1965). Afterwards, he studied Soils and Hydrological Sciences in France. His research interests focus on soil reclamation in the Atacama Desert, irrigation with saline waters in native Andean communities, arsenic and boron contamination in water and crops, preservation of water quality in mining areas. In 1995, he received a National Government Award for his contributions to the use of waters in the arid zone of northern Chile. Now, he is Professor of Environmental Geochemistry at UCN and a consultant on water quality management for the National Environmental Commission, a Chilean government agency, and for several mining companies. Carlos Salazar is a Civil Engineer from the Universidad de Chile (1981). He joined the Chilean National Water Resources Service (Direccion General de Aguas), where he is principally involved in water resources development and global management as the Head of the Studies and Planning Department. He has special interests in groundwaters and wetland systems in arid areas.
The origin of brines and salts in Chilean salars: a hydrochemical review Francßois Risacher a,*, Hugo Alonso b, Carlos Salazar c b
a CNRS-IRD, Centre de Ge´ochimie de la Surface, 1 rue Blessig, 67000 Strasbourg, France Universidad Cato´lica del Norte, Departamento de Quı´mica, Angamos 0610, Antofagasta, Chile c Direccio´n General de Aguas, Ministerio de Obras Pu´blicas, Morande´ 248, Santiago, Chile
Received 26 June 2002; accepted 20 February 2003
Abstract Northern Chile is characterized by a succession of north – south-trending ranges and basins occupied by numerous saline lakes and salt crusts, collectively called salars. Fossil salt crusts are found to the west in the extremely arid Central Valley, while active salars receiving permanent inflows fill many intravolcanic basins to the east in the semiarid Cordillera. Sea salts and desert dust are blown eastward over the Cordillera, where they constitute an appreciable fraction of the solute load of very dilute waters (salt content < 0.1 g/l). The weathering of volcanic rocks contributes most components to inflow waters with salt content ranging from 0.1 to 0.6 g/l. However, the average salt content of all inflows is much higher: about 3.2 g/l. Chemical composition, Cl/Br ratio, and 18O – 2H isotope contents point to the mixing of very dilute meteoric waters with present lake brines for the origin of saline inflows. Ancient gypsum in deep sedimentary formations seems to be the only evaporitic mineral recycled in present salars. Saline lakes and subsurface brines are under steady-state regime. The average residence time of conservative components ranges from a few years to some thousands years, which indicates a permanent leakage of the brines through bottom sediments. The infiltrating brines are recycled in the hydrologic system where they mix with dilute meteoric waters. High heat flow is the likely driving force that moves the deep waters in this magmatic arc region. Active Chilean salars cannot be considered as terminal lakes nor, strictly speaking, as closed basin lakes. Almost all incoming salts leave the basin and are transported elsewhere. Moreover, the dissolution of fossil salt crusts in some active salars also carries away important fluxes of components in percolating brines. Evaporative concentration of inflow waters leads to sulfate-rich or calcium-rich, near-neutral brines. Alkaline brines are almost completely lacking. The alkalinity/calcium ratio of inflow waters is lowered by the oxidation of native sulfur (reducing alkalinity) and the deposition of eolian gypsum (increasing Ca concentration). Theoretically, SO4-rich inflow waters and their derived SO4-rich brines should be found in the intravolcanic basins of the Cordillera because of the ubiquity of native sulfur, while Ca-rich brines should prevail in sedimentary basins where Ca-rich minerals are abundant. This relation is perfectly observed in the salar de Atacama, the largest in Chile. However, several salars located within the volcanic Cordillera belong to the Ca-rich group. Inflows and brines may have acquired their Ca-rich composition in Pleistocene time when their drainage basins were mainly sedimentary. Later on, recent lava flows and ignimbrites covered the sedimentary formations. Underground waters may have kept their early sedimentary signature by continuous
* Corresponding author. Tel.: +33-3-90240409; fax: +33-3-88367235. E-mail addresses: [email protected] (F. Risacher), [email protected] (H. Alonso), [email protected] (C. Salazar). 0012-8252/$ - see front matter D 2003 Elsevier Science B.V. All rights reserved. doi:10.1016/S0012-8252(03)00037-0
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F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
recycling. However, the weathering of volcanic rocks tend to slowly shift the water compositions from the Ca-rich to the SO4-rich type. D 2003 Elsevier Science B.V. All rights reserved. Keywords: Northern Chile; Hydrochemistry; Closed basin; Salar; Salt recycling; Brine evolution
1. Introduction The Central Andes in Chile, Bolivia and Argentina contains a large number of closed basins whose central depression is occupied by saline lakes and salt crusts (salars). Active research has mainly been focused on the world’s largest salt crust: the salar de Uyuni in the Bolivian Altiplano (Ericksen et al., 1978; Rettig et al., 1980; Risacher and Fritz, 1991a, 2000; Fornari et al., 2001; Baker et al., 2001). Information on other small salars are not easily available. They can be found with difficulty in many technical reports, theses and communications at local symposia and bear mostly on economic geology and water resources of the basins. Nevertheless, Stoertz and Ericksen (1974) and Chong (1984) described in detail the features of the salar crusts of northern Chile, but gave little hydrochemical information. Risacher and Fritz (1991b) studied the geochemistry of most of the small intravolcanic salars of the southern Bolivian Altiplano. The origin of solutes in a few Chilean salars has been investigated by Alpers and Whittemore (1990), Spiro and Chong (1996), Alonso and Risacher (1996), Garce´s (2000a,b) and Carmona et al. (2000). Other studies of Chilean salars deal with paleohydrography and paleoclimatology (Grosjean, 1994; Valero-Garce´s et al., 1996; Bao et al., 1999; Grosjean et al., 2001), sedimentology (Valero-Garce´s et al., 1999) and water quality (Caceres et al., 1992; Aguirre and Clavejo, 2000). In this paper, we present a review of the hydrogeochemistry of almost all active salars of northern Chile: 52 saline lakes and salt crusts are distributed along a 1000 km north – south transect in the Andes Cordillera. The drainage basin of an important river and its tributaries (Rio Lauca), feeding salar de Coipasa in the Bolivian Altiplano, has been included. We do not undertake a detailed geological, mineralogical, sedimentological and paleolimnological study of each salar, but to review as a whole the major
hydrogeochemical trends, focusing on the origin of salts and brine evolution. This approach is possible because all these salars belong to the same geographical, geological and climatological environment in spite of their scattering over a 200,000 km2 area. The large number of basins and water samples smoothes local irregularities and provides a global view of their chemical trends. Fig. 1 presents the study area and the location of all salars. Each basin or salar is identified in the text by a number (from Fig. 1) bracketed after its name or abbreviation. Table 1 gives the detailed location and the physiographic and climatic features of the basins. All minerals cited in the text are listed with their formula in Table 2.
2. Analytical data This study is largely based on the analyses of 354 inflow waters and 227 brackish lake waters and brines of Chilean salars (Risacher et al., 1999). However, bromine is not included in this report. Inasmuch as the Cl/Br ratio is an important clue to the origin of solutes, we present in Appendix A the previously unreported Br analyses together with Cl concentration. Both components were analysed by ion chromatography (Dionex 4000i). A discussion of the hydrogeochemical setting of each salar can be found in Risacher et al. (1999). For the giant salar de Atacama, we rely on the comprehensive studies of Moraga et al. (1974, 206 analyses) and Ide (1978, 389 analyses).
3. Regional geology Northern Chile is characterized by contractional structures and intensive block faulting due to the subduction of the Nazca Plate below the South American Plate. A succession of north – south-trending
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
251
Fig. 1. Map of northern Chile and topographic sections showing the main morphostructural units and the location of salars. The elongated and thin formation in the Pre-Andean Depression, between salar de Atacama and Domeyko Range, is the Cordillera de la Sal.
252
Table 1 Physiographic and climatic features of closed basins in northern Chile Code
LAGUNAS COTACOTANI LAGO CHUNGARA RIO LAUCA SALAR DE SURIRE SALAR DE PINTADOS LAGUNA LAGUNILLA SALAR DEL HUASCO SALAR DE COPOSA SALAR DE MICHINCHA SALAR DE ALCONCHA SALAR DE CARCOTE SALAR DE ´N ASCOTA SALAR DE ATACAMA SALAR DE TARA LAGUNA HELADA SALAR DE AGUAS CALIENTES 1 SALAR DE PUJSA SALAR DE LOYOQUES LAGUNA TRINCHERA
COT
No. (Fig. 1)
Latitude (South)
Longitude (West)
Altitude Salar (m)
Alt max (m)
1
18j11V08
69j13V11
4495
6342
CHR
2
18j14V48
69j09V13
4530
LAA SUR
3 4
18j27V00 18j50V03
69j14V54 69j02V08
PIN
5
20j31V38
LGU
6
HCO
Temperature mean (jC)
Prec (mm/ year)
Evap (mm/ year)
TDS min (mg/l)
TDS max (mg/l)
1.9
379
1070
299
2002
SO4
119
6
6
6342
1.9
338
1230
47
1633
SO4
273
22.5
22.5
3892 4260
6063 5780
4.2 2.7
370 250
1200 1280
108 108
784 285,000
SO4 SO4
2374 574
– 144
– 9.5
69j40V17
980
5785
18.5
75
1630
199
365,621
SO4
17,150
377
0
19j55V56
68j50V40
3900
5190
4.6
150
1490
267
SO4
129
7
20j18V18
68j50V22
3778
5190
5
150
1260
108
113,093
SO4
1572
51
2.5
COP
8
20j38V12
68j39V33
3730
5220
5
150
1300
119
330,671
SO4
1116
85
5
MIC
9
20j58V48
68j33V09
4125
5407
3.5
200
1620
94
62,662
SO4
282
2.5
0.6
ALC
10
21j03V07
68j29V14
4250
5407
3.5
200
1620
116
99,361
SO4
128
3.8
0.75
CAR
11
21j22V36
68j20V55
3690
6176
5.8
125
1630
88
335,536
SO4 – Ca
561
108
ASC
12
21j33V01
68j16V58
3716
6023
5.8
125
1630
89
119,853
SO4 – Ca
1757
243
ATA
13
23j25V05
68j15V57
2300
6233
14
160
1800
243
339,719
SO4 – Ca
18,100
3000
TAR
14
23j01V59
67j16V30
4400
5816
0
150
1500
287
176,486
SO4
2035
48
HEL
15
23j05V58
67j08V15
4300
5745
0
180
1500
118
308,546
SO4
221
AC1
16
23j07V22
67j24V38
4280
5370
1
150
1500
1359
122,890
Ca
281
15
2.5
PSA
17
23j12V26
67j30V35
4500
6046
1
150
1500
239
57,578
SO4
634
18
5
LOY
18
23j14V36
67j16V55
4150
5370
1
150
1500
163
242,456
Ca
676
80
5
TRI
19
23j24V01
67j23V51
4290
5060
0
200
1500
10,830
10,830
SO4
59
1276.0
Evol. path
Area basin (km2)
Area Salar (km2)
0.2
5.8
0.4
Area water (km2)
0.15
3.5 18 12.6 14 5.8
0.3
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
Basin
MUE
20
23j26V08
67j26V28
4295
5109
0
200
1500
30,916
97,153
SO4
41
AC2
21
23j28V51
67j33V11
4200
6046
1
150
1500
2530
13,656
Ca
1168
LEJ MIS
22 23
23j29V44 23j43V19
67j41V28 67j45V58
4325 4120
5924 5910
1 2
150 180
1500 1500
1273 217
70,723 5308
SO4 SO4
193 303
MIN
24
23j45V46
67j47V30
4120
5910
2
180
1500
3702
10,912
SO4
LAC TUY
25 26
23j51V15 23j56V22
67j24V24 67j35V23
4250 4010
5852 5852
1 1
200 180
1500 1500
636 6572
24,433 131,024
SO4 SO4
AC3
27
23j59V16
67j41V33
3950
5910
1
150
1500
2491
25,150
CPR
28
23j57V25
67j47V28
3950
5204
1
150
1500
6589
IMI
29
24j11V06
68j47V00
2949
3995
10
40
2000
PUN
30
24j35V38
69j00V00
2945
6739
10
150
AC4
31
25j00V16
68j37V15
3665
6739
2
LAZ
32
25j04V18
68j30V33
4250
5697
PAJ
33
25j08V33
68j48V50
3537
GOR
34
25j24V07
68j40V27
IGN
35
25j29V39
AZU
36
AMA
0.03
134
0.03
9
1.9 13.4
1.9 13.4
1.6
1.6
306 245
16.2 2.9
2.2 2.9
SO4 – Ca
476
46
2.5
221,804
Ca
137
27
0.9
565
90,008
Ca
189
2000
196
233,367
Ca
4263
180
1630
851
341,759
SO4
656
19.5
1
180
1630
3965
158,854
SO4
393
7.7
5488
5
115
1350
11,728
246,674
Ca
1984
104
1.4
3950
5467
1
140
1000
2846
296,737
SO4
324
27
0.5
68j37V09
4250
5100
2
140
1000
4043
97,091
SO4
25j28V36
68j48V30
3580
5488
3
120
1100
548
323,473
Ca
214
37
25j32V56
68j49V59
3558
5488
2
120
1100
7656
196,672
Ca
863
23
0.04
AGI
38
25j48V06
68j53V24
3320
5058
2
100
1100
177,044
334,882
Ca
589
71
0.0002
ISL
39
25j44V50
68j38V00
3950
5761
0
130
1000
6229
329,693
SO4
858
152
27.5
37.5
9.8 250
0.16 0.1
2
3.5
0.7
0.002
3.3
0.02
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
LAGUNA CHIVATO MUERTO SALAR DE AGUAS CALIENTES 2 LAGUNA LEJIA LAGUNA MISCANTI LAGUNA ˜ IQUE MIN SALAR DE LACO LAGUNA TUYAJTO SALAR DE AGUAS CALIENTES 3 SALAR DE CAPUR SALAR DE IMILAC SALAR DE PUNTA NEGRA SALAR DE AGUAS CALIENTES 4 LAGUNA DE LA AZUFRERA SALAR DE PAJONALES SALAR DE GORBEA SALAR IGNORADO SALAR DE LA AZUFRERA SALAR DE AGUA AMARGA SALAR DE AGUILAR SALAR DE LA ISLA
2.4
253
(continued on next page)
254
Table 1 (continued) Evap (mm/ year)
TDS max (mg/l)
Evol. path
0
140
1000
8907
333,942
SO4
676
40
1.2
5664 5012
2 2
130 100
1000 1100
8277 1677
129,707 318,744
Ca SO4 – Ca
867 293
29 6.7
0.4 0
3494
5373
3
120
1100
3288
20,430
SO4
400
69j08V51
3370
6127
4
125
1200
85
326,745
26j20V12
68j44V49
4150
5857
2
140
1000
2908
82,327
46
26j18V09
68j36V24
4250
6080
2
140
1000
2134
JIL
47
26j23V10
68j39V41
4150
6127
2
140
1000
BAY
48
26j23V09
68j35V02
4250
6127
2
140
WHE
49
26j41V35
68j37V32
4220
6146
1
ESC
50
26j38V51
68j31V21
4353
5918
LAV MAR
51 52
26j52V39 26j55V25
68j27V49 69j04V31
4350 3760
FRA
53
27j27V47
69j13V03
4110
Code
No. (Fig. 1)
Latitude (South)
Longitude (West)
Altitude Salar (m)
Alt max (m)
SALAR DE LAS PARINAS SALAR GRANDE SALAR DE INFIELES SALAR DE LA LAGUNA SALAR DE PEDERNALES SALAR DE PIEDRA PARADA LAGUNAS BRAVAS LAGUNAS DEL JILGUERO LAGUNA DEL BAYO SALAR DE WHEELWRIGHT LAGUNA ESCONDIDA LAGUNA VERDE SALAR DE MARICUNGA LAGUNA DEL NEGRO FRANCISCO
PAR
40
25j49V33
68j30V17
3987
5761
GRA INF
41 42
25j59V52 25j58V32
68j42V00 69j03V26
3950 3520
LGN
43
26j12V29
68j58V29
PED
44
26j13V42
PIE
45
BRA
Temperature mean (jC)
TDS min (mg/l)
SO4 – Ca
Area basin (km2)
Area Salar (km2)
0.55
Area water (km2)
0.4
3620
335
0.6
SO4
388
28
0.2
121,494
SO4
545
10
2925
19,592
SO4
119
3.4
1.5
1000
2432
4785
SO4
221
1.2
1.2
140
1000
1466
158,739
Ca
466
6.3
4
1
140
1000
2318
32,315
SO4
194
3.8
3.5
6893 6893
1 2
170 120
1000 1100
935 144
174,581 331,453
SO4 SO4 – Ca
6052
1
200
1000
130
327,885
SO4
1075 3045 933
15 145 24.8
10
15 6 24.8
Climatic data are taken from the Balance Hidrico de Chile (DGA, 1987). The potential evaporation (Evap) measured in class A evaporimeters has been corrected by a factor 0.65. Alt max is the highest elevation of the drainage basin. TDS min and max are the lowest and the highest salt contents of the waters and brines, respectively, found in each basin. Basin area includes drainage area and salar area. Salar area includes lakes, ponds and dry surfaces (crusts, playas).
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
Prec (mm/ year)
Basin
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293 Table 2 Minerals cited in the text and their formula Amorpous silica Antarcticite Bischofite Calcite Gypsum Halite Mirabilite Natron Thenardite Trona Ulexite
SiO2 CaCl26H2O MgCl26H2O CaCO3 CaSO42H2O NaCl Na2SO410H2O Na2CO310H2O Na2SO4 Na2CO3NaHCO32H2O NaCaB5O98H2O
ranges and basins alternate in the orogenic belt of the Central Andes (Fig. 1). The main morphostructural units are from West to East: (1) the Coast Range; (2) the Central Depression; (3) the Precordillera; (4) the Pre-Andean Depression; (5) the Western Cordillera; (6) the Altiplano; (7) the Eastern Cordillera. The Coast Range is a moderately elevated mountain chain (1500 m as an average) steeply dipping westward in the Pacific Ocean. Marine sedimentary rocks and volcanic andesitic rocks of Mesozoic age make up most of the range. The Central Depression (or Central Valley), at an altitude of 800 to 1400 m, is a thick fill of detrital and lacustrine sediments (gravel, sand, silt, clay) of middle Tertiary to Holocene age, often covered with salt crusts remnants of ancient salars. Roughly between 22j and 25jS latitude, the Central Depression is known as the Atacama Desert. The famous Chilean nitrate deposits are located in this depression where they impregnate many surficial formations (Ericksen, 1981). On the east margin of the valley, large alluvial fan systems make the transition with the Precordillera Range. The Domeyko Range, at an elevation of 3500– 4000 m, is the most conspicuous part of the Precordillera because it is clearly separated from the Eastern Cordillera by the Pre-Andean Depression. The Precordillera Range is made up chiefly of Paleozoic and Mesozoic sedimen-
255
tary, metamorphic and igneous rocks. The large copper ore deposits of northern Chile are located in the Precordillera. The Domeyko Range also contains continental evaporite formations on its eastern fringe. The Pre-Andean Depression, also called the salar de Atacama Depression (not to be confused with the Atacama Desert), is a large intramontane basin, at 2500 m of elevation, filled with Tertiary to Holocene clastic and evaporitic sediments of continental origin. It contains the largest evaporitic basin in Chile: the salar de Atacama ([13], 3000 km2). Another important salar, Punta Negra ([30], 250 km2), occupies the southern edge of the basin. A prominent ridge of anticlinal folds and domes of gypsum and rock salt borders the salar de Atacama on its western fringe: the Cordillera de la Sal. It corresponds to an ancient salar of Tertiary age deformed as a consequence of tectonic events during the Cenozoic (Dingman, 1962, 1967). The Western Cordillera, of Miocene to Holocene age, is an elevated plateau (4000 m) consisting of rhyolitic ignimbrites and covered by numerous andesitic stratovolcanoes reaching 6500 m in elevation. Many of these volcanoes contain native sulfur deposits, some of them of economic interest. The volcanoes often delineate interior drainage basins occupied by saline lakes and salt crusts. Most of the salars included in this study belong to this unit. The Bolivian Altiplano is a major intramontane basin separating the Western Cordillera from the Eastern Cordillera above 22jS latitude. Its central trough is occupied by the giant salar de Uyuni (10,000 km2). The volcanic rocks of the Western Cordillera are in contrast to the sedimentary rocks that dominate in all other units, from the Coast Range to the Precordillera.
4. Climate Climatic features of northern Chile are summarized in Table 3 (DGA, 1987). The potential evaporation
Table 3 Climatic features of northern Chile
Precipitation Potential Evaporation Mean temperature
Coast Range
Central Valley
Precordillera
Pre-Andean Depression
Western Cordillera
< 10 1000 15 – 18
< 10 1500 18 – 20
10 – 25 1500 8 – 12
25 – 50 1500 12
50 – 300 600 – 1200 0
Precipitation and evaporation are in millimeters/year and temperature in jC.
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F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
given in Tables 1 and 3 is derived from measurements in class A evaporimeters (DGA, 1987). Pan evaporation is usually greater than the actual evaporation from lake surface estimated by other methods (hydrologic balance, energy budget). This is due to the differences in size, heat distribution and water movement between large lakes and small and shallow evaporation pans. An empirical pan coefficient is generally applied to correct evaporimeter measurements. The many weaknesses of this approach has been discussed by Miller (1977, p. 266). Unfortunately, at most places, they are the only available data. Reasonable values for pan coefficients may be extrapolated from studies by Jones (1965), Van Denburgh (1975), Yonts et al. (1973), Smith (1979) and Stanhill (2002). Most pan coefficients range from 0.6 to 0.8. In northern Chile, Grilli et al. (1986) selected a single 0.65 correction term, which will be used in this study. A salient feature of the climate is the increase in precipitation from west to east across this region. The coast range and the Central Depression are among the driest deserts of the world, in contrast with the semiarid Western Cordillera where precipitation reaches 300 mm/year. In the whole area, however, the potential evaporation greatly exceeds the precipitation, which is a basic condition for the establishment of evaporitic environments (Hardie and Eugster, 1970; Eugster and Hardie, 1978). In the northern area, 10% to 20% of the precipitation between 4000 and 5000 m occur as snows. In contrast, in the southern area, snow constitutes as much as 50% to 80% of the precipitation at the same elevation (Vuille, 1996). The mean annual temperature decreases from north to south, which allows the snow blanket to persist for a longer time in the southern basins. The dominant wind direction is from ocean to cordillera. Thus, sea salts and saline dust from the Central Valley are transported eastward over the Highlands by the wind.
5. Main outlines of Chilean salars 5.1. Salar morphology A comprehensive description and many photographs of the surficial features of active and fossil
salars, from the Coast Range to the Highlands, is given by Stoertz and Ericksen (1974). Active Chilean salars of the Western Cordillera resemble those described in the Bolivian Altiplano by Risacher and Fritz (1991b). Three main types are found in northern Chile as follows. (1) Saline lakes are permanent bodies of saline waters. The water area occupies the whole salar area (see Table 1). They range in depth from some decimeters to some meters. Mirabilite often precipitates in winter because its solubility is markedly reduced by low temperatures. (2) Playas are the most common type of salar in the Western Cordillera of northern Chile, as they are in the Bolivian Altiplano. Shallow ( < 20 cm) and ephemeral saline pools occupy variable areas of wide mudflats across the central depression of the basin. Some meters below the playa surface, a confined aquifer is filled with an interstitial brine, whose capillary draw precipitates a variety of salts within the sediments (gypsum, mirabilite, halite, ulexite). (3) Salt crusts of two kinds are found in this region. Active salt crusts are presently forming from evaporating brines that have reached saturation with respect to gypsum, mirabilite or halite. These crusts are of low thickness, rarely above 50 cm, and are found in the semiarid Western Cordillera. Fossil salt crusts are located in more arid areas, from the Coast Range to the western fringe of the Western Cordillera. These crusts are of similar mineralogy to the active crusts, but are much thicker. The halite crust of salar de Atacama is several hundred meters thick (Bevacqua, 1991; Bobst et al., 2001). However, fossil crusts are more commonly of metric to decametric size. A 10-mthick halite crust of Neogene age was described in the Coast Range by Chong et al. (1999). Deep dissolution pits (halite, Pajonales [33]), outcrops above the salar level (thenardite, Gorbea [34]), tilted layers (gypsum, Pedernales [44]) are evidence of fossil, nonactive, salt crusts. Many salars are combinations of these types, especially salt crusts and playas. Generally, salars in the Coast Range, in the Central Valley and in the Precordillera are fossil salt crusts, while those in the Western Cordillera are predominantly playa lakes. Mixed salars are found in the Central Depression and in the western fringe of the Western Cordillera,
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293
close to the Precordillera. In this study, we have incorporated the largest salar of the Central Valley, salar de Pintados, because its underground waters are actively fed by inflows originating in the Western Cordillera. However, the salar itself is a large fossil salt crust.
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et al., 1984; Hastenrath and Kutzbach, 1985; Sylvestre et al., 1999; Baker et al., 2001; Fornari et al., 2001). It is likely that the neighbouring Chilean basins were also filled with saline lakes during the same period. Fossil salt crusts stem from the desiccation of such old saline lakes.
5.2. Inflows 6. Inputs Four types of inflows feed Chilean salars are as follows: (1) Focused discharge through springs from the drainage area with an observable flow rate generally between 0.1 and 100 l/s. Many springs emerge far above the salar surface. (2) Diffuse seeps percolate slowly to the salar surface. Their flow rate is not noticeable, but their lateral extension may reach tens or even hundreds of meters along the salar shore. Total discharge through diffuse seeps may be significant but is impossible to quantify. The slow flow rate may induce an incipient evaporation of the waters at the seepage face. (3) Rivers are generally of short length and low discharge, generally below 500 l/s. They often infiltrate in colluvium before reaching the salars, where they eventually discharge through springs and seeps. (4) Underground waters fill aquifers in the drainage basins. They discharge in the central trough of the basins through springs and seeps. It was often possible to sample these waters in exploration and exploitation wells drilled by mining companies. 5.3. Former lakes Since Miocene time, the climate have remained arid or semiarid in northern Chile (Ko¨tt et al., 1995; Pueyo et al., 2001). Many basins have been occupied by saline lakes (Stoertz and Ericksen, 1974; Geyh et al., 1999). Late Neogene lacustrine formations are known in the Central Valley (Sa´ez et al., 1999; May et al., 1999) and in the Lauca Basin in the Western Cordillera (Gaupp et al., 1999). Grosjean (1994) and Grosjean et al. (1995) described a late Pleistocene lacustrine phase in Laguna Lejia [22]. Several large and deep saline lakes have repeatedly covered the Bolivian Altiplano during Pleistocene time (Servant and Fontes, 1978; Lavenu
6.1. Atmospheric contribution Two fresh snows and seven late-lying snows were sampled and analysed for major and minor components. The meltwater from one old snow has a salt content as high as 23 mg/l due essentially to Ca and SO4, while the salt content of the six others ranges from 1.0 to 6.6 mg/l (mean: 3.6 mg/l). In comparison, meltwaters from fresh snows have salt contents of 3.1 and 3.5 mg/l (mean: 3.3 mg/l). Therefore, six out of seven old snows did not receive significant dry fallout after their deposition, which suggests that these snows were probably sampled shortly after their fall. Only one late-lying snow was markedly contaminated by eolian gypsum. It is therefore possible to obtain a broad estimation of the flux of each component brought by snowfalls in the basins by multiplying the volume of precipitation over each basin by the average concentration of the meltwaters. 6.2. Surface and groundwater inputs On the other hand, the flux of components entering permanent lakes through springs, seeps, rivers and underground discharges can also be estimated. The hydrologic balance of a saline lake without surface or underground outflow is given by (Hutchinson, 1957; Langbein, 1961): Vinflows ¼ S ðE PÞ
ð1Þ
where Vinflows = basin discharge to the lake (m3/year), S = surface of the lake (m2), E = actual evaporation (m/ year) and P = precipitation (m/year). Water loss from Chilean saline lakes through infiltration is generally small compared to basin discharge (see discussion below on salt balance in
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Section 8.2). The flux of components brought in the lake by all inflow waters is obtained by multiplying their average concentration by the volume of inflows. The uncertainty on the basin discharge depends on that of the climatic parameters E and P and on the unweighted average concentration of inflow waters, except when the lake is obviously fed by one major input. Figs. 2 and 3 compare the fluxes of 11 components brought by inflows to 12 permanent lakes to their fluxes deposited by snows on their drainage basins. Lakes are ordered according to the increasing average salt content of their inflows. Two groups can be
distinguished: lakes fed by dilute inflows whose salt content is lower than 500 mg/l (CHR [2], HEL [15], COT [1], LGU [6]) and lakes fed by inflows of higher salt content: 500 –6500 mg/l (MIS [23], LEJ [22], FRA [53], BRA [46], ESC [50], LAV [51], MIN [24], TUY [26]). These two groups of waters differ also significantly in relative ionic concentrations. 6.3. Lakes fed by dilute inflows 6.3.1. Major components in dilute inflows The flux of major components brought by dilute inflows in lakes CHR [2], HEL [15], COT [1], LGU
Fig. 2. Fluxes of major components deposited by snows on the drainage basins (stars) of 12 permanent lakes compared to the fluxes brought into the lakes by inflow waters (squares). Focusing our attention on the SO4 diagram, we can see that the flux of SO4 brought by dilute inflows is only slightly more elevated than that brought by snows on the drainage basins (for LGU [6], the proportion is even reversed). In contrast, concentrated inflows bring in the lakes significantly more SO4 than snows. The other major components show similar behavior although it is not so easily noticeable because of the vertical logarithmic scale.
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[6], are only 1 to 14 times higher than those deposited by the snows (Fig. 2). In LGU basin, all fluxes are even slightly higher in snows than in inflow waters. This means that for LGU, as well as for other lakes fed by dilute waters, atmospheric deposition could represent a major, or the major, contribution to the lakes. This seems particularly clear for SO4 in very dilute waters, which is probably derived from eolian sulfates (gypsum and thenardite) and oxidized elemental sulfur eroded from the top of the volcanoes. Dissolved components in very dilute groundwaters stem from rock alteration and atmospheric inputs that have been brought from the land surface to the groundwater by percolating waters. The flux of atmo-
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pheric inputs deposited on the land surface should then be lower than the flux brought from the basin by inflow waters. A lower flux from the basin could be due to the removal of atmospheric components in soils. We must also lay stress on the large uncertainty on flux estimates. 6.3.2. Minor components in dilute inflows With the only exception of As in Laguna Helada (HEL [15]), the flux of minor components (Br, I, NO3, As) deposited by the snows are 1 to 70 times higher than those brought into the lakes by dilute inflows (Fig. 3: CHR [2], HEL [15], COT [1], LGU [6]), which suggests that most of the minor compo-
Fig. 3. Fluxes of minor components deposited by snows on the drainage basins (stars) of 12 permanent lakes compared to the fluxes brought into the lakes by inflow waters (squares). Except for LGU, most of Si is brought in the lakes by inflow waters. In contrast, the fluxes of Br, I, NO3 and As deposited by the snows are higher than those brought in the lakes by dilute inflows. Not so for concentrated inflows, which bring in the lakes more minor components than snows, with the exception of NO3.
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nents in dilute waters are probably brought by precipitation. Br and I may be derived from sea salts, and NO3 and I may be derived from surficial formations of the Central Valley. The removal of atmospheric components in soils and plants could be very efficient for minor components such as NO3 and I. Mineral deposits and geothermal areas are the primary sources of arsenic in arid regions (Smedley and Kinniburgh, 2002). Owing to the high aridity, As is not leached from the surficial formations and constitutes a noticeable component of desert dust, which makes it a major health hazard in northern Chile (Smith et al., 1998; Queirolo et al., 2000a,b). Its concentration in meltwaters ranges from 0.3 to 68 Ag/l (mean: 18 Ag/l), compared to the world baseline concentration in rain and snow of 0.01 to 0.5 Ag/l (Smedley and Kinniburgh, 2002). In addition, copper smelting plants in northern Chile permanently release large quantities of As in the atmosphere (Gidhagen et al., 2002). As for silica, in all basins but LGU [6], the flux brought by inflow waters is markedly higher (10 to 75 times) than the flux deposited by the snows. Silica is likely to originate from the dissolution of volcanic glass. 6.4. Lakes fed by concentrated inflows Except for NO3, the fluxes of all components deposited by the snows are small in comparison with those brought by concentrated inflows: approximately 3 to 800 times lower. NO3 seems to be the only component, among those studied, that originates in all waters mostly from atmospheric deposition. Br, I, Si and As in these waters must come from some additional source. In the above discussion, we considered only the flux of components brought in the basins by snowfall. However, during the much longer dry season, dry fallout is also deposited on the drainage basins. Therefore, the fluxes of all components brought by atmospheric inputs are very likely to be higher than those estimated only from the snows. This strengthens the conclusion that the lowest the concentration of an inflow water is, the higher is the contribution of atmospheric salts to its salt load. Atmospheric salt is the main source of dissolved components in very dilute waters. As concentration increases, other inputs become progressively more significant.
7. Origin of salts in inflow waters 7.1. Salt content of inflow waters The most conspicuous feature of inflow waters in closed basins of northern Chile is their very broad range of concentration: from 47 to 79,740 mg/l with an average of 3230 mg/l. Northern inflows are globally less saline (average = 2.4 g/l) than southern inflows (average = 4.9 g/l). In comparison, White et al. (1963, 1980) reported salt contents of underground waters in andesitic to rhyolitic aquifers of North America ranging from 70 to 473 mg/l. Alteration modeling of a standard andesite (code Kindisp, Made et al., 1994) leads to a maximum concentration of 340 mg/l of dissolved solids. Thus, the alteration of volcanic rocks does not seem to be the main controlling factor of the high concentration of many inflow waters. The concentration of most components in inflow waters increases with the total salt content (Figs. 4 and 5). The correlation line between component concentration and salt content shows similar slopes for inflows and lakes (Na, Cl, Li>1 and K, Mg, SO4, Ca < 1). The concentration increase in lake waters is easily explained by evaporative concentration, a mechanism that cannot significantly account for most inflow waters. The rate of evaporation of groundwater is much lower than in an open water and decreases drastically with depth (Ripple et al., 1972; Hellwig, 1973; Miller, 1977, p. 304; Yechieli and Wood, 2002). One may reasonably neglect evaporation of deep groundwater within the unsaturated zone. However, shallow groundwater ( < 3 to 5 m below surface) is subject to interstitial evaporation, especially when the capillary fringe reaches the surface. A conspicuous feature of this process is the deposition of efflorescent salts on soils and plants around the salars. Therefore, evaporative concentration may increase the salt content of shallow groundwater at the margin of the salars. However, many deep waters and springs with high discharges also have elevated salt content that cannot be related to direct evaporative concentration. The simplest hypothesis is that dilute waters have mixed with underground brines of composition similar to the present lake brines. The proportion of most components of the deep brines would be nearly conserved by dilution with freshwaters, which could explain the similar slopes observed in Fig. 4.
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Fig. 4. Concentration of major components in inflow waters and lake brines of Chilean salars as a function of total dissolved solids (TDS). Note the overall correspondence between the slopes of the correlation lines of inflows and lakes.
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Fig. 5. Concentration of minor components in inflow waters and lake brines of Chilean salars as a function of total dissolved solids (TDS).
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7.2. Mixing with underground brines Sodium + calcium ( 2) + magnesium ( 2) is plotted against chloride in Fig. 6. Dilute inflows (Cl < 10 mmol/l) present a large scatter in values, while saline inflows (Cl>10 mmol/l) are gathered along the isoconcentration line Cl = Na + 2Ca + 2Mg. Rock alteration is likely to be the main control of the composition of dilute inflows. North American waters from andesitic rocks have compositions similar to the dilute waters. This diagram is alike to those representing any common evaporative system as illustrated in the inset diagram showing a similar relation for evaporated lake waters. However, all waters represented in the main diagram are inflow waters: groundwaters, springs, seeps and rivers. Only seeps and shallow groundwaters may have undergone some evaporation just before they emerge. Nevertheless, seeps and springs present the same behavior, which suggests that the evaporative trend shown in the main diagram could have been inherited from ancient evaporated brines buried below the volcanic formations. Alternatively, the shift to the isoconcentration
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line could be due to the dissolution of a mixture of sodium chloride with calcium and/or magnesium chloride in dilute underground waters. Ancient evaporites containing halite and gypsum are present below the volcanic formations. However, calcium and magnesium chloride are very soluble salts (antarcticite: 1650 g/kg [H2O] and bischofite: 1170 g/kg [H2O]) known to form either in extremely cold environments or in very concentrated brines. They are unlikely to be present in depth at relatively elevated temperature. A more probable explanation is the mixing of dilute waters with Na –Ca – Mg/Cl underground brines similar to those present in saline lakes as shown in the inset diagram in Fig. 6. 7.3. Mixing with subsurface brines Mixing of underground water with brines at the margin of a salar follows the Ghyben – Herzberg principle, which defines the geometry of a freshwater lens of low density lying on a large saltwater body (Heath, 1983). The typical case is a fresh groundwater lens in an island. The Ghyben– Herzberg principle has
Fig. 6. Sodium + calcium ( 2) + magnesium ( 2) vs. chloride concentration for inflow waters (main diagram) and for lake brines (inset diagram) of Chilean salars. North American waters from andesitic rocks are included for comparison. Both diagrams show the same overall pattern, which suggests that inflow waters stem from the mixing of freshwaters with brines similar to those present in the salars.
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received little attention in the case of playas and saline lakes (Yechieli and Wood, 2002). It is difficult to apply to salars in northern Chile. Firstly, the overall pattern is inverted: brine lenses overlie freshwater bodies. This is a basically unstable situation: the denser brine lens tends to move downward. The Ghyben – Herzberg mixing zone slopes away from a saline lake (see Yechieli and Wood, 2002), in contrast with a freshwater lens overlying a saline layer (see Heath, 1983, p. 68). It is not easy to extrapolate very much in depth the interface between fresh and saltwater below a saline lake. The geometry of the lens can be reasonably described only close to the surface, near the zone of discharge. Secondly, the Ghyben – Herzberg principle applies only to an isotropic aquifer, which is not the case of volcanic formations with fracture porosity. Water and brine may be forced to move along unpredictable pathways. Thirdly, thermal heating below the basin may considerably disturb the interface between the water bodies. Anyway, brine mixing occurs at the margin of the salars and could be partially responsible of the elevated salt content of the seeps. However, many springs emerging far above the
salar surface also have high salt content. Thus, mixing at the margins is not likely to be the main process of the high concentration of inflow waters. 7.4. Geochemistry of bromide Chloride-to-bromide ratios shown in Fig. 7 also suggest that the composition of saline inflow waters (Cl>10 mmol/l) results from the mixing of snow meltwaters with brines similar to those found in the present lakes. Two mixing curves between average snow meltwater and lake brines are drawn: curve A corresponds to the lake brine of lowest Cl/Br ratio and curve B to that of highest Cl/Br ratio. Any other mixing curve is constrained between these two end member curves. All saline inflows plot in this mixing area. By contrast, several dilute inflows (Cl < 10 mmol/l) plot outside the mixing area. Br content of halite associated to Tertiary gypsum diapirs in the neighbouring Bolivian Altiplano is around 2 ppm (Risacher and Fritz, 2000). Similar low values (1– 2 ppm) have been found by Pueyo et al. (2001) in neogene salt rocks of northern Chile. The Cl/Br ratio
Fig. 7. Cl/Br ratio as a function of Cl for snows, inflow waters and lake brines. Two end member mixing curves between snow meltwater and lake brines are drawn: Curve A corresponds to the lowest Cl/Br ratio (alc-9) and curve B corresponds to the highest ratio (isl-12). All other possible mixing curves are located between these two curves. Curve C shows the Cl/Br ratio of the solution produced by dissolving halite in snow meltwater. Curve C is well outside the field of saline inflows.
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obtained by dissolving such low Br content halite in pure water is figured by an horizontal dashed line far above all inflow points. Curve C represents the Cl/Br ratio of the solution obtained by dissolving halite in average snow meltwater. The halite endpoint may be put anywhere on the horizontal dashed line for Cl values between 355 and 46,500 mg/l, which is the highest Cl content measured in saline inflows. The dissolution curve C if well above all saline inflow points. Thus, the high salt content of inflows cannot stem from the dissolution of ancient salt rock overlapped by volcanic formations. 7.5. Stable isotopes The d2H and d18O values for 89 inflow waters, 6 lakes (Chaffaut, 1998) and 26 snows (Pen˜a, 1989) are shown in Fig. 8. The meteoric water line for northern Chile is taken from Aravena et al. (1999). The lakes were sampled in the central part of the studied area ( f 23jS latitude) and the snows in the northernmost sector ( f 18j). Most of the snows show a marked displacement to the right of the meteoric line. According to Pen˜a (1989), the shift is due to sublimation and evaporation of the snow blanket during the day. Night
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condensation does not counterbalance the evaporation in this arid region. Moser and Stichler (1980) carried out experimental evaporation of snow samples which led to a similar shift due to d2H and d18O increase. Inflow waters are clearly distributed along a line joining the lightest isotopic values to the heavily enriched ones observed in the lakes. The isotopic values of inflow waters follow geographic and hydrologic trends. (1) Northern inflows ( < 24j latitude) have lower isotopic values than southern inflows; only two northern inflows have high isotopic values similar to those found in the south. (2) Seeps are more enriched in d2H and d18O. The shift of inflow waters to the right of the meteoric line is due to the conjunction of several processes reflected by the two correlations. Snow prevails over rain and remains for a longer time in the south than in the north of the study area (see Section 4). As a result, the evaporation of the snows and their related isotopic enrichment observed in the north by Pen˜a (1989) is likely to be significantly enhanced in the south. Seep waters are subjected to evaporation by capillary draw before they emerge and by direct evaporation as they percolate very slowly to the land surface. There is no surprise if these waters have
Fig. 8. 18O and 2H contents in snows, inflow waters, and some lake brines of Chilean salars. The isotopic composition of the most concentrated inflow of salar de Surire [4] that results from mixing of groundwater with brine is shown. The size of lake symbols (open circles) is not related to their salt content.
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Fig. 9. Ionic activity products of calcite, amorphous silica and gypsum as a function of total dissolved solids (TDS). Square brackets refer to activities. Horizontal lines show the solubility product at 0 and 25 jC. A water is supersaturated with respect to a phase when its representative point plots above the solubility product line. Below the line, the water is undersaturated with respect to the phase. Many inflow waters are oversaturated, or close to saturation, with respect to calcite and amorphous silica. These two minerals seem to exert a main control on Si, Ca and carbonates concentration of most inflow waters. As TDS increases, inflow waters tend to reach saturation with respect to gypsum. Eventually, gypsum is saturated, or very close to saturation, in 3 springs and 10 seeps, which suggests gypsum dissolution, at least in spring waters.
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high d2H and d18O values. In contrast, underground waters, springs and rivers (of short length in the study area) do not undergo significant evaporation before sampling, but also present a significative isotopic enrichment, particularly in the south. The higher salt content of many of the southern waters associated with heavier isotopic contents suggests rather a mixing process than an evaporative concentration. Dilute and concentrated endpoints are snow meltwaters and brines, respectively, similar to those found in present saline lakes. The relation between salt content and isotopic enrichment suggested in Fig. 8 is partially obscured by the respective salt contents of the two endpoints. For example, the water resulting from the mixing of a little volume of hypersaline brine with a large volume of dilute water has an isotopic composition close to that of the dilute water. In contrast, the salt content of the mixed water may be markedly higher than that of the dilute water. The volume of saline water is negligible, but not the mass of dissolved salts. A typical example is that of the most saline inflow (17.7 g/l) to the salar de Surire [4]. The spring emerges within the salar after passing through a thick section of brine-filled sediments. The high salt content
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is obviously not due to evaporative concentration of the spring water, but to the mixing with a little volume of hypersaline brine, which is confirmed by the d2H and d18O values plotting almost on the meteoric water line. 7.6. Mineral dissolution Fig. 9 shows the ionic activity product of three important minerals involved in the salt load of inflow waters and in their evaporative evolution: calcite, gypsum and amorphous silica. Solute and water activities are calculated with the ion interaction model (Pitzer, 1979; Harvie et al., 1984). H4SiO4 activity and amorphous silica solubility are taken from Marshall (1980), Marshall and Warakomski (1980) and Marshall and Chen (1982). The software used is EQL/ EVP (Risacher and Clement, 2001 and complete database references herein). Most of the inflow waters are saturated, or close to saturation, with respect to calcite and amorphous silica. Therefore, the solubility of amorphous silica controls the concentration of silica in inflow waters. Calcite is very likely an alteration mineral that controls, at least partially, the concentration of Ca
Fig. 10. Temperature (jC) of springs and underground waters and mean annual air temperature as a function of elevation above sea level. Waters that plot above the curve of mean annual air temperature + 5 jC present a thermal influence.
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and HCO3 in inflow waters. Most inflow waters are undersaturated with respect to gypsum. However, as concentration increases, gypsum becomes close to saturation in 10 seeps and 3 springs. Concerning the seeps, it is possible that their Ca and SO4 concentration has increased by evaporation. However, this cannot be the case for springs. For these waters, the dissolution of gypsum is the simplest way to reach saturation. There are no other common minerals that
could supply sufficient quantities of Ca and SO4 to reach saturation with gypsum. 7.7. Thermal waters Northern Chile is a region of high heat flow due to recently active volcanism (Springer and Fo¨rster, 1998). Geothermal systems, such as that of El Tatio (outside the study area), have been considered for
Fig. 11. Lithium and boron content of springs and underground waters [in wt.% of total dissolved solids (TDS)] as a function of temperature. No correlation with temperature is observed.
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energy production (Lahsen, 1988). Thermal influence can only be estimated in springs and underground waters. The temperature of seeps and rivers is controlled by the atmosphere. The temperature of springs and underground waters is compared in Fig. 10 to the mean annual air temperature (Grilli et al., 1986; DGA, 1987). If we reasonably assume that a water whose temperature is at least 5 jC higher than the mean annual air temperature presents a thermal influence, then thermal heating has affected 78% of all sampled springs and underground waters. However, the thermal influence is moderate: only 10 waters have temperatures between 40 and 53 jC, which are the highest values observed in the study area. No correlation between temperature and solute concentration
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has been found. Fig. 11 shows the concentration of Li and B (in wt.% of the salt content) as a function of water temperature. Li and B are generally enriched in thermal waters (Berthold and Baker, 1976; White et al., 1976). Nevertheless, in northern Chile, such an enrichment is not observed. Therefore, hydrothermal alteration does not seem to be a major process of solute acquisition. Fig. 12 shows the spatial distribution of inflows differentiated according to the difference between their temperature and that of the atmosphere (Twater –Tatm; seeps and rivers are not considered). In the northern and central area (18 – 24jS latitude), both thermal and nonthermal waters are intermingled. In the south (24 – 27jS latitude), however, a clear separation can be observed. Nonthermal springs are located westward close to the boundary between the sedimentary Precordillera and the volcanic Western Cordillera. In contrast, heated inflows are found preferentially eastward well within the volcanic Cordillera. This observation is in good agreement with the increase of heat flow from the Coast Range to the active magmatic arc of the Western Cordillera reported by Springer and Fo¨rster (1998).
8. Origin of deep brines 8.1. Recycling of ancient salars
Fig. 12. Distribution of thermal springs and underground waters in salar basins of northern Chile. Note the exaggerated horizontal scale. In the south, thermal inflows are more abundant eastward within the Cordillera. A similar relation may also be suspected in the northwestern part of the study area.
Volcanic formations in northern Chile range in age from Miocene to Pleistocene. During this entire time span, the climate remained arid or semiarid (Ko¨tt et al., 1995; Pueyo et al., 2001). Many salars must have occupied ancient intravolcanic basins. Subsequently, lava flows and ignimbrite sheets covered these saline deposits. The buried brines may have mixed with dilute meteoric inflows to produce brackish and saline underground waters. However, this straightforward explanation for the origin of deep brines comes up against two difficulties. First, the brines must have been preserved during the hot volcanic events. Elevated temperatures may vaporize the lake waters. Only underground brines protected by overlying sediments could have been preserved. Secondly, these entrapped brines must have been rapidly exhausted as they discharge with fresh inflows. The time of exhaustion depends of the size of the buried salars
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and of their leaching rate by dilute waters. We may intend a rough calculation of the exhaustion time based on present salars. Their wide range of morphological and chemical types make it difficult to define an average salar. On the other hand, extreme values of salar size and water composition may not be very significative. Therefore, we have selected a broad range of values for each parameter that covers a large number of salars of different type: salar surface = 10 to 150 km2; lake surface = 1 to 15 km2; salt content of inflows = 0.5 to 5 g/l; salt content of brines = 20 to 200 g/l; evaporation = 1.33 m/year; precipitation = 0.16 m/year. The only parameter that cannot be easily estimated is the volume of underground brines. We will assume a thickness of 100 m of brinefilled sediments with a porosity of 40%. The mass of dissolved salts in deep brines of a buried salar can be estimated by multiplying the salar surface by the sediment thickness, the porosity and the brine concentration. The flux of dissolved salts brought each year in a present salar is obtained by multiplying the volume of inflows [see above, Eq. (1)] by their salt content. By dividing these two values, we obtain an exhaustion time ranging from 100 years (a small salar rapidly leached, generating concentrated inflow waters) to 1.7 million years (a large salar slowly leached, generating dilute waters). Small- and medium-sized salars have been leached in a short time as compared with the age of volcanic formations of Central Andes (Miocene to Holocene). Only large buried salars could still provide some salts in present inflow waters. 8.2. Brine leakage from present salars Brine leakage from present salars is another source of salts in underground waters. The concentration of a conservative component (Li, Br, Cl before halite saturation) in a saline lake is controlled by the infiltration rate of the lake brine through bottom sediments. Water leakage as a controlling factor of lake chemistry has already been stressed by Carmouze and Pedro (1977), Alderman (1983), Sanford and Wood (1991) and Dutkiewicz et al. (2000). If a lakebed is completely impermeable, then the concentration of a conservative component, and consequently the salt content of the lake, will continuously increase, which in turn reduces the evaporation rate of
the lake brine. A steady state is attained at very high concentration when the equilibrium water vapor pressure exerted by the solution equals the mean relative humidity of the atmosphere (Kinsman, 1976). In contrast, if infiltration occurs through the lakebed, even at a very slow rate, then a steady state is attained when the input flux balances the leakage flux of the conservative component. The difficulty is to determine if the steady state is attained. A recent climatic, geologic or anthropic event may have abruptly modified the lake composition, which could be undetectably reequilibrating through time with the new climatic or geologic conditions. One way is to compare the total mass of the component in the lake to its annual input flux. The ratio between total mass and annual input has the dimension of a time. If this time is substantially lower than the age of the last disturbing event, then it is referred to as the residence time of the component in the lake. A short residence time indicates a high infiltration rate. The total mass of a component in the lake is: Mlake ¼ Slake Hlake Clake
ð2Þ
where Slake is the surface of the lake, Hlake is its mean depth and Clake is the average concentration of the selected conservative component in the lake. The annual input flux is [see Eq. (1)]: Minflows ¼ Slake ðE PÞ Cinflows
ð3Þ
where E is the actual evaporation, P is the precipitation and Cinflows is the average concentration of the conservative component in inflows. The ratio between the total mass in the lake and the annual input flux is: Mlake =Minflows ¼ Hlake =ðE PÞ Clake =Cinflows
ð4Þ
Table 4 presents this ratio for Cl in some characteristic Chilean saline lakes. Most lakes belong to the Alconcha [10] – Lagunilla [6] types with ratios ranging from a few years to some hundreds. Laguna Helada [15] has the highest estimated ratio: 3370 years. The last drastic climatic change occurred 10,000 years ago when large lakes occupying the basins dried up (see Section 5.3). It seems therefore reasonable to assume that all these ratios represent the residence time of Cl and that Andean saline lakes are under steady-state
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Table 4 Rough estimates of residence times and infiltration rates of some typical north Chilean saline lakes Lake surface (km2) Mean lake depth (m) Mean inflow concentration (Cl: mmol/l) Lake concentration (Cl: mmol/l) Residence time [years, Eq. (4)] Infiltration rate [mm/year, Eq. (5)] Loss through infiltration (106 l/year) Basin discharge [106 l/year, Eq. (1)] Percent infiltration relative to basin discharge
HEL [15]
FRA [53]
ALC [10]
MIS [23]
LGU [6]
5.8 0.50 0.64 5050 3370 0.15 0.86 6800 0.01
14.3 0.20 2.0 4810 411 0.5 7.0 16,700 0.04
0.75 0.20 0.80 1030 220 0.90 0.68 880 0.08
13.4 4.5 1.8 34 73 62 830 15,700 5.3
0.15 0.15 21 161 1 150 23 180 13
Cl is the conservative component used for the calculations. Lake levels may change during the year. Cl concentration corresponds to the lake size given in this table. FRA surface is different from that given in Table 1. Actually, FRA is constituted of two separated lakes: one dilute lake that acts as a preconcentration pool and a hypersaline lake which is considered here. MIS mean depth is taken from Valero-Garce´s et al. (1996) and FRA mean depth from Behnke (1987). In both cases, we have halved the maximum depth given in these studies to obtain an approximate mean value inasmuch as the surface of both lakes reported by these authors was similar to that observed during our field trips. All other mean depths were estimated by ourselves. The main uncertainty is the mean Cl concentration of inflows, difficult to estimate. Underground discharges and ephemeral streams during the wet season are not taken into account.
regime as well as most subsurface brines. The salt balance for a conservative component (Cl, Li, Br) is then given by: Vinflows Cinflows ¼ Vleakage Clake
ð5Þ
where C is the average concentration of the component and V is the annual flux of solutions. Eq. (5) allows to estimate leakage rates from saline lakes. Table 4 gives infiltration rates of characteristic Chilean saline lakes. The volume of lake brine lost by infiltration ranges from 0.01% to 12.8% of the basin discharge. Therefore, it is generally justified to neglect the leakage in the water balance of a saline lake although it is the main control of the salt balance of the lake. Almost all dissolved salts entering the lake are lost by leakage. Only low soluble salts (calcite, gypsum, Mg silicates) are removed from the solution and stored in lake sediments. Their mass is very small compared to the total input flux to the lake. 8.3. Dissolution of present salt crusts Most of the incoming solutes are lost by infiltration, which precludes the formation of thick salt crusts. However, some salars, particularly those located close to the Precordillera, contain several meter thick halite crusts filled with an interstitial
brine (Imilac [29], Pajonales [33], Gorbea [34], Agua Amarga [37], Aguilar [38], Pedernales [44]). These thick crusts are quite different from those of centimetric to decimetric size forming by evaporation in other salars (see Section 5.1). Numerous deep brine ponds suggest that the thick crusts are undergoing active dissolution. They do not seem to be presently forming. Under former wetter climate, deep salt lakes occupied the basins (see Section 5.3). An abrupt drying-up due to a climatic change induced the rapid evaporation of the lake and the precipitation of a thick salt crust. The residual lake is often limited to a subsurface, intracrustal brine saturated with respect to halite. The intracrustal brine infiltrates through bottom sediments removing large quantities of dissolved salts. In contrast, inputs to the salar are fresh or brackish waters largely undersaturated with respect to halite, which tends to dilute the intracrustal brine. In order to maintain saturation with respect to halite, the crust must dissolve. As long as the salt crust remains in the salar, the flux of infiltrating salts is higher than the flux of incoming salts, which induces a slow dissolution process responsible for the numerous ponds pitting the crust. A new steady-state regime will be established only when the whole crust has vanished. In summary, ancient salt crusts in present Andean salars are not likely to be in equilibrium with their climatic and hydrologic environment. They are an additional
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source of salts in underground waters close to the Precordillera. 8.4. The salt cycle in Chilean salars Salts do not accumulate in most Andean salars. The salt cycle is only delayed. The delay is the residence time of the dissolved components. Most incoming salts only pass through the saline formations, where they undergo a concentration process and are further recycled in the hydrologic system. Infiltrating brines are not necessarily recycled in the same basin. The intense and complex fracturing of volcanic rocks could favor deep circulation from one basin to the others. This is a continuous mechanism that accounts well for the high salt content of inflow waters over long time spans. Thermal convection is the likely driving force that mixes, moves and heats the underground waters in this area of active volcanism. There is no need of a large salt reservoir below the volcanic structures. Nevertheless, giant salt bodies are present in depth. Alonso et al. (1991) estimate at 10,000 km3 the volume of deep halite in Central Andes. The Bolivian Altiplano is dotted with gypsum diapirs. The Cordillera de la Sal in northern Chile plunges to the east below ignimbrites and lava flows. Gypsum recycling through volcanic activity in the Western Cordillera has been reported by Risacher and Alonso (2001). Presently, only gypsum seems to supply part of Ca and SO4 to the salars. Owing to its high solubility, the top of halite bodies formerly in contact with underground waters could have been leached a long time ago and redeposited in the giant salars of Uyuni and Atacama. Only low soluble gypsum is still in contact with subsurface waters and constitutes a present source of Ca and SO4.
9. Brine evolution 9.1. The model Waters undergoing evaporative concentration precipitate a sequence of minerals in order of increasing solubility. Hardie and Eugster (1970) and Eugster and Hardie (1978) introduced the basic concept of chemical divide to account for the evolution of solute concentration when a water evaporates. In brief sum-
mary, when a mineral precipitates, the concentration of all its components cannot increase simultaneously in the solution because the ionic activity product of the mineral must be kept constant and equal to its solubility product. For example, in the case of a mineral containing one anion and one cation, if the anion activity increases in the solution, then the cation activity must decrease. The solution becomes enriched in some components and depleted in others, according to the precipitated minerals and to the concentration ratio of their components at the start of the precipitation. Two approaches may be used to predict the fate of an inflow water upon evaporation. The quantitative approach calculates step by step the chemical composition of the solution and the sequence and quantities of minerals produced during evaporation. In this study, we use the EQL/EVP code (Risacher and Clement, 2001) based on the ion interaction model (Pitzer, 1979; Harvie et al., 1984). The qualitative approach only intends to predict which pathway a water will follow, without caring about the precise chemical composition of the solution. The main difficulty lies in that many analyses found in the literature, and all of this study, report total alkalinity (Alk) as a routinely measured parameter. The first chemical divide produced by the early precipitation of calcite would apparently necessitate to calculate the respective concentrations of CO32 and HCO3. The first model established by Hardie and Eugster (1970) was grounded on carbonate speciation, neglecting ion association and using the Debye – Hu¨ckel equation for ion activities. The [Ca2 +] vs. [CO32 ] concentration relation is a rather complex expression because [HCO3] and [OH] must also be taken into account. In addition, each analysis must be recomputed so that the water is in equilibrium with the atmospheric pCO2. It is possible to update these calculations by using the ion interaction model and the EQL/EVP code. However, this would be a useless repetition of the quantitative approach. A much simpler and rigorous treatment directly based on the alkalinity concept has been developed by Al-Droubi et al., (1980) for the first chemical divides in which carbonate and silicate species are involved. We will summarize this approach which does not require any computer code (see also Risacher and Fritz, 1991b).
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The total alkalinity of a natural water can be expressed as (Dickson, 1981; Stumm and Morgan, 1996): Alk ¼ 2½CO2 3 þ ½HCO3 þ ½OH þ ½BðOHÞ4
½Hþ
ð6Þ
where square brackets refer to total concentrations in mol/l or mmol/l. This formulation corresponds to the operational definition of alkalinity, but is difficult to handle because all terms are interdependent. By combining Eq. (6) with the electro-neutrality equation, we obtain the alternative definition of alkalinity: Alk ¼ ½Naþ þ ½Kþ þ 2½Ca2þ þ 2½Mg2þ ½Cl 2½SO2 4
ð7Þ
Alkalinity is the difference between the sum of the conjugate cations of the strong bases and the sum of the conjugate anions of the strong acids (Stumm and Morgan, 1996, p. 165). Calcite is the first mineral to precipitate from most natural solutions undergoing evaporative concentration. We may rewrite Eq. (7) as: Alk 2½Ca2þ ¼ ½Naþ þ ½Kþ þ 2½Mg2þ ½Cl 2½SO2 4
ð8Þ
As long as calcite is the only precipitate, the concentration of each component of the right-hand side of Eq. (8) increases linearly with the concentration factor F of the evaporating solution: ½Naþ ¼ F ½Naþ 0
ð9Þ
where F ¼ ðH2 OÞ0 =ðH2 OÞ
ð10Þ
(H2O) is the mole number of water in the solution. Subscript ‘‘0’’ refers to the initial solution at the start of evaporation. Eq. (8) becomes: Alk 2½Ca2þ ¼ F ð½Naþ 0 þ ½Kþ 0 þ 2½Mg2þ 0 ½Cl 0 2½SO2 4 0 Þ
ð11Þ
which leads to: Alk 2½Ca2þ ¼ F ðAlk0 2½Ca2þ 0 Þ
ð12Þ
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If at the start of evaporation Alk0>2[Ca2 +]0, then the right-hand side of Eq. (12) is positive and increases linearly with the concentration factor. Thus, as evaporation proceeds, the difference between alkalinity and calcium concentration will continuously increase on behalf of alkalinity. Conversely, if Alk0 < 2[Ca2 +]0, then calcium concentration will increasingly predominate over alkalinity. Inasmuch as carbonate species are the main contributors to alkalinity in most natural waters, the behavior of alkalinity reflects that of carbonates. Therefore, the first chemical divide of a water composition upon evaporation can be very easily and rigorously determined by examining the simple ratio of the analytical concentrations Alk0/ 2[Ca2 +]0 of the initial inflow water. Mg carbonate or Mg silicates also precipitate at an early stage of water evolution. The effect of coprecipitation of calcite and Mg salts on the evolution of an evaporating water may be determined by rewriting Eq. (7) as: Alk 2½Ca2þ 2½Mg2þ ¼ ½Naþ þ ½Kþ ½Cl 2½SO2 4
ð13Þ
which leads straightforwardly to: Alk 2½Ca2þ 2½Mg2þ ¼ F ðAlk0 2½Ca2þ 0 2½Mg2þ 0 Þ
ð14Þ
If Alk0>2[Ca2 +]0 + 2[Mg2 +]0, then the difference between Alk and 2[Ca2 +] + 2[Mg2 +] is positive and continuously increasing. The evaporating water evolves into an alkaline brine depleted in Ca and Mg and enriched in carbonates. Conversely, if Alk0 < 2[Ca2 +]0 + 2[Mg2 +]0, then the water follows a neutral path depleted in carbonate. If Alk0>2[Ca2 +]0 and Alk0 < 2[Ca2 +] 0 + 2[Mg2 +]0, then the evaporating water first follows the alkaline path as long as calcite is the sole precipitate and reverts to the neutral path when Mg carbonate or Mg silicates reach saturation. It must be emphasized that either Mg carbonate or Mg silicates have the same effect upon the chemical divide. Fig. 13 shows some initial evolutionary paths of north Chilean inflow waters in a diagram Alk vs. 2[Ca]. Waters following the neutral path become progressively enriched in Ca and, in most cases, reach saturation with gypsum (CaSO 42H2O). Gypsum
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Fig. 13. Examples of evaporation paths of north Chilean inflow waters illustrating the qualitative approach (analyses in Risacher et al., 1999: 1 = asc-14; 2 = mic-1; 3 = ped-9; 4 = sur-26; 5 = sur-7). Solid circles are directly plotted from water analyses. Quantitative evaporation paths are computed with the EQL/EVP evaporation program based on the ion interaction model and a rigorous treatment of the aqueous species. At the beginning of evaporation, all waters remain undersaturated with respect to calcite. Ca and Alk behave conservatively and each water follows a straight line path whose equation is 2[Ca]/Alk = 2[Ca]0/Alk0. 2[Ca]0 and Alk0 are the coordinates of the starting points (solid circles 1 to 5). The precipitation of calcite causes an abrupt deviation of the linear evaporation path depending on which side of the iso-concentration line Alk = 2[Ca] the starting point stands. The precipitation of magnesite in path 5 reverses the alkaline path to the neutral path because Alk0 < [Ca]0 + 2[Mg]0 in inflow water 5. The precipitation of Mg salts also occurs in paths 3 and 4, but does not reverse the alkaline paths inasmuch as Alk0>2[Ca]0 + 2[Mg]0 in these waters. The pH of the initial solution may be modified by CO2 exchange with the atmosphere or by CO2 uptake by aquatic plants. The result is a displacement of the calcite appearance point along the initial conservative straight line, which does not change the evolutionary trend.
induces a new divide of the neutral path producing either Ca-rich/SO4-poor waters or Ca-poor/SO4-rich waters, according to their Ca/SO4 ratio at the beginning of gypsum saturation (not at the beginning of evaporation). The whole discussion is summarized in the flow sheet of Fig. 14 which accounts for most natural waters (modified from Hardie and Eugster, 1970). A few minerals play a prominent part in the water evolution. The outcome of the processes is the generation of three main groups of brines: alkaline (Na/ HCO3 –CO3 –Cl), sulfate-rich (Na/SO4 – Cl) and calcium-rich (Na –Ca/Cl) brines. The flow sheet of Fig. 14 defines four pathways: alkaline (I – IA); sulfatealkaline (I – IIA –III); sulfate-neutral (II – III) and calcic (II – IV). Pathways I– IIA –IV and II –(Na/CO3 – Cl) have not been observed in nature.
Actually, Na and Cl are the main components of most brines, but they do not induce a chemical divide before halite saturation, which occurs at high concentration once the major evolutionary trends of the brine have already been defined. Throughout the first crucial steps of evaporative concentration, Na and Cl behave as conservative components. For the sake of brevity, waters, where SO4>Ca, will be referred to as SO4-rich, and those where Ca>SO4 will be referred to as Ca-rich. The composition of a brine is strongly dependent on that of the initial dilute water. A small variation of the ratio of the critical components in the dilute water {Alk0/2[Ca]0, Alk0/(2[Ca]0 + 2[Mg]0)} may cause drastic changes in the final brine composition. In turn, the composition of dilute inflow waters is mainly inherited from rock alteration. Therefore, there is an
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Fig. 14. Flow diagram for evaporative concentration of dilute waters. Path numbers are those of Hardie and Eugster (1970): alkaline (I – IA); sulfate-alkaline (I – IIA – III); sulfate-neutral (II – III) and calcic (II – IV). The dashed line shows a rarely observed direct path from the Mg salts divide to the sulfate-rich brine group, bypassing the gypsum divide. It has no particular incidence for water evolution and is included in our discussion in the sulfate-alkaline path I + IIA + III.
overall relation between lithology and brine composition. Each broad type of lithology is reflected in the composition of the dilute water which is further amplified by evaporative concentration. In a pioneering study, Garrels and Mackenzie (1967) first threw light on the (nonobvious) relation between igneous rocks and alkaline brines. High-purity igneous rocks are almost devoid of anionic components (Cl , SO42 ). The only available anion that can balance the cations derived from the weathering of pure igneous rocks is HCO3, which stems from atmospheric CO2 and is brought by the rains as carbonic acid. However, igneous rocks often contain mineralization or Cl – SO4-rich fluid inclusions which lowers the proportion of HCO3 in the weathering solution. Moreover, sulfide minerals and native sulfur not only
provide SO42 but also acidify the solution through oxidation reactions such as: Sj þ 3=2O2 þ H2 O ! 2Hþ þ SO2 4
ð15Þ
The alkalinity of the solution may be drastically reduced. Therefore, inflow waters draining moderately mineralized igneous rocks tend to follow the sulfate alkaline evaporative pathway, while highly mineralized igneous rocks produce waters following the sulfate-neutral pathway. Extreme cases of very intense acidification leading to acid-sulfate brines have been reported in Australia (McArthur et al., 1991; Long et al., 1992) and Chile (Risacher et al., 2002). Sedimentary rocks are often characterized by high Ca (limestone, marl, dolomite, gypsum). Waters
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in contact with such rocks generally tend to generate calcic brines. 9.2. Discrepancies between model and observed brine evolution The model relating lithology, inflow waters and brine composition sets the geochemical and hydrochemical baseline of evaporative concentration. It assumes direct evaporation of the inflow water as it would occur in an artificial pool protected from any contamination. Actually, several processes may disturb this scenario. The evaporating water often interacts with eolian dust, sediments or salts, which modifies its composition and may change its theoretical evolutionary path. For example, the behavior of potassium is mostly controlled by exchange reactions on clay minerals (Eugster and Jones, 1979). The content of SO4 may be lowered by bacterial reduction or freezing out of Na sulfate salts (Eugster and Hardie, 1978). Gac (1980) carried out evaporation experiments of the main tributary of Lake Chad with and without allowing contact with atmospheric dust. The behavior of Mg was very different in each experiment. In Bolivian Andes, native sulfur is eroded from the top of the volcanoes and deposited on the drainage basins where its oxidation acidifies inflow waters enough to reverse their initial alkaline evolutionary path to the sulfatealkaline or sulfate-neutral path (Risacher and Fritz, 1991b). These interactions are most efficient at an early stage of evaporative evolution, when inflow waters are still dilute. Brackish and saline inflows are much less sensitive to them. Another disturbing mechanism is the mixing of several inflows with different evolutionary paths. Mixing in large basins of complex lithology is a more likely scenario than in small monolithologic basins. Leakage through bottom sediments has also been reported to significantly modify the brine evolution (Sanford and Wood, 1991). All these well-documented processes take place after discharge of inflow waters in the basin, while they undergo evaporative concentration. In this paper, we describe a (apparently) new discrepancy between lithology and evolutionary paths of inflow waters. Several Chilean salars located in an exclusive volcanic environment are fed by inflow waters following the calcic path typical of sedimentary rocks. Their brines do not reflect the geological context.
9.3. Water evolution in Chilean salars 9.3.1. The calcite divide As shown in Fig. 9, most inflow waters are close to saturation, or oversaturated, with respect to calcite. The relations Alk vs. 2Ca and Alk vs. 2Ca + 2Mg for inflow waters are shown in Fig. 15. At the start of evaporation (upper diagram), 238 inflow waters (67%) should follow the initial neutral pathway (II), while 117 (33%) should follow the opposite initial alkaline pathway (I). The precipitation of Mg salts (lower diagram), probably Mg smectites as in Bolivian salars (Badaut and Risacher, 1983), theoretically shifts 65 out of 117 waters from the initial alkaline pathway to the sulfate-alkaline pathway (I – IIA), remaining 52 waters (15%) that should end up as carbonate brines (I – IA). In contrast to the prediction, in the whole study area, only one small brackish pond located in the drainage basin of laguna Helada [15] is actually of alkaline composition (HEL [15] in the lower diagram). All other 226 lakes and ponds contain low-alkaline or neutral waters of the Na/Cl– SO4 or Na – Ca/Cl groups including the main laguna Helada itself. The discrepancy has already been observed and discussed in the neighbouring Bolivian salars and explained by acidification due to the oxidation of eolian sulfur in the drainage basins (Risacher and Fritz, 1991b). It is very likely that the same process occurs in northern Chile, where many volcanoes also contain similar native sulfur deposits. However, the proportion of inflow waters that should end up as carbonate brines in Bolivian Andes is much higher than in northern Chile: 52% and 15%, respectively. Moreover, 6 Bolivian salars out of 30 are actually alkaline. Therefore, an additional process must be involved to explain the lower proportion of alkaline waters and brines in Chilean Andes. One of the most likely is dry fallout of desert dust from the Central Valley on the Chilean Andes. The importance of atmospheric inputs in the composition of very dilute waters has already been emphasized (see Section 6). Gypsum is one of the main components of desert and atmospheric dust. Its dissolution increases the concentration of Ca of the inflow waters without modifying their alkalinity, which may shift the water evolution from the alkaline path to the neutral or to the sulfate-alkaline paths.
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Fig. 15. 2Ca and 2Ca + 2Mg concentration vs. alkalinity for inflow waters in Chilean salars. Lake brines values are also plotted on the 2Ca + 2Mg diagram. To avoid redundancy, waters with Alk < 2Ca are not plotted in the lower diagram. Inset flow diagrams show the evaporative concentration paths followed by the group of waters differentiated by the isoconcentration lines. There is only one lake brine (HEL) in the alkaline-rich field although 52 inflow waters should end up after evaporation in the alkaline-rich field.
9.3.2. The gypsum divide The upper diagram of Fig. 16 shows the Ca and SO4 contents of all inflows and lake waters. Lake waters are differentiated according to their Ca/SO4 ratios: 143 are sulfate-rich and 82 Ca-rich waters. In contrast, inflow waters are differentiated according to their evolutionary path calculated with the EQL/EVP code (Risacher and Clement, 2001). Many inflow waters that end up as sulfate brines have initial concentrations higher in Ca than in SO4. The early
precipitation of calcite reverses this ratio allowing the solution to become sulfate-rich. Two typical evolutionary paths are drawn: one (line A) leading to a Na/Cl– SO4 brine (but starting in the Ca>SO4 area) and the other (line B) ending in a Na – Ca / Cl brine. Owing to the large number of inflows (354), it is not possible to show the theoretical evolutionary path of each of them and to compare it to the actual composition of the lake or pond it feeds. Actual evolutionary paths after the gypsum divide are in
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Fig. 16. SO4 and SO4 + Alk/2 concentration vs. Ca for inflow waters and lake brines in Chilean salars. Inset flow diagrams show the evaporative concentration paths followed by each group of waters. The evolutionary path of each inflow water has been determined with the EQL/EVP evaporation program. In contrast, lakes are simply differentiated according to their Ca/SO4 ratio. Many waters following the SO4-rich path are actually in the Ca>SO4 field. The early precipitation of calcite lowers the Ca concentration and reverses the Ca/SO4 ratio. Line A shows such an evolutionary path. Line B shows a typical Ca-rich path.
general agreement with the calculated paths. We could not detect reversion from one path to the other, as observed after the Mg salts divide when waters reverse from the alkaline path to the sulfatealkaline path due to eolian sulfur. The gypsum divide
occurs at rather high Ca and SO4 concentrations once their ratio can no longer be easily modified. Nevertheless, the mixing of brackish waters of different Ca/SO4 ratio may produce unpredictable end brines.
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In total, 243 inflow waters follow either the sulfate-alkaline or the sulfate-neutral path which ends up as SO4-rich solutions in 143 lakes and ponds, while 60 inflows follow the calcic path leading to Ca-rich solutions feeding 82 lakes and ponds (Fig. 16). In addition to the 243 inflows following the sulfate paths, 52 inflows that theoretically should follow the alkaline path up to carbonate brines also end up as sulfate-rich solutions after reverting to the sulfate-alkaline path (Fig. 15). All the 225 lakes and ponds (143 + 82) are distributed in 52 salars. Table 1 shows to which chemical group each salar belongs. Approximately two-thirds of the salars (34 out of 52, SO4 in Table 1) contain only brackish and saline waters high in SO4 and low in Ca. Twelve salars belong to the Ca-enriched group (Ca) and seven contain both Ca-rich and SO4-rich solutions. Several salars of the Ca-rich group have a few SO4-rich brackish ponds. However, their most concentrated brines are Ca-rich. The symmetric situation is not observed in SO4-rich salars, where Ca-rich brines are completely lacking.
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Salar de Atacama, which is the largest in Chile, contains both Ca-rich and SO4-rich brines within the halite nucleus (Risacher and Alonso, 1996). The western part of the nucleus, close to the Cordillera de Domeyko where sedimentary rocks are predominant, is filled with Ca-rich brines, while the eastern part close to volcanic formations of the Cordillera is filled with SO4-rich brines. The Ca vs. SO4 diagram shows very clearly the gypsum divide and the relation between lithology and brine composition (Fig. 17; data from Moraga et al., 1974; Ide, 1978). As can be seen in the upper diagram of Fig. 16, it is not possible to use the initial Ca/SO4 ratio of the inflow water to predict which evaporative path will be followed after the gypsum divide. In contrast, as shown in the lower diagram, Ca vs. SO4 + Alk/2, the isoconcentration line Ca = SO4 + Alk/2 separates perfectly the two groups of inflow waters. Those with initial Ca>SO4 + Alk/2 follow the Ca-rich path, and those with SO4 + Alk/2>Ca follow the SO4-rich path. This is obvious if Ca>SO4 + Alk/2: whatever the mass of CaCO3 and CaSO4 precipitated, Ca will remain in
Fig. 17. SO4 vs. Ca diagram illustrating the gypsum divide in salar de Atacama brines. (These data are not included in Fig. 16.) Points are distributed along a tilted T. The ascending branch (+) corresponds to waters undersaturated with respect to gypsum. Waters saturated with gypsum plot along the approximate gypsum saturation line and are divided into two groups: SO4-rich brines (circles) above the isoconcentration line Ca = SO4, and Ca-rich brines (squares) below. The geograpical distribution of the three groups of solutions is shown in the Atacama outline map. Analyses are from Moraga et al. (1974) for the SO4 – Ca diagram and from Ide (1978) for brine distribution in the salt nucleus. Note that the approximate gypsum saturation line has no rigorous chemical meaning, inasmuch as the true saturation line can only be drawn in an activity diagram.
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Table 5 Theoretical evaporative paths and final brine composition predicted from the initial composition of inflow waters Concentrations in inflow water
Evaporative path
Brine group
Alk>2Ca and Alk>2Ca + 2Mg Alk>2Ca and Alk < 2Ca + 2Mg Alk < 2Ca and SO4 + Alk/2>Ca Alk < 2Ca and SO4 + Alk/2 < Ca
alkaline (I – IA)
Na/CO3 – Cl
sulfate-alkaline (I – IIA – III) sulfate-neutral (II – III) calcic (II – IV)
Na/SO4 – Cl Na/SO4 – Cl Na – Ca/Cl
Path numbers as in Fig. 14. Concentrations are in mol/l or mmol/l.
excess. However, this rule cannot be rigorously demonstrated in the case SO4 + Alk/2>Ca. If the precipitation of calcite is not occurring and if Ca>SO4, then the Ca-rich path could be followed. The most common way to prevent, or to delay, the precipitation of calcite is to increase the content of CO2 gas in the
water. This happens in confined aquifers where the water is not in contact with the atmosphere and therefore does not evaporate. In evaporitic environments, the CO2 dissolved in waters is close to equilibrium with that of the atmosphere. Thus, the ratio (SO4 + Alk/2)/Ca of the initial water may be used with good confidence to predict which of the Ca-rich or SO4-rich path will be followed by an evaporating water after the gypsum divide. Table 5 summarizes the relations between the initial composition of inflow waters and their evaporative path. 9.3.3. Relation between evolutionary path and temperature We have plotted in Fig. 18 the ratio (SO4 + Alk/2)/ Ca of all springs and underground waters against the difference between water temperature and mean annual air temperature (seeps and rivers are not considered). Positive Y-axis values (log scale) correspond to the SO4-rich path and negative values
Fig. 18. Diagram showing the relation between temperature and the two main evolutionary paths SO4-rich and Ca-rich for springs and underground waters. Waters are differentiated according to their salt content. Most hot inflow waters follow the SO4-rich path. Note that the region labeled the ‘‘SO4-rich path’’ include the waters that have reversed their theoretical evolutionary path from alkaline-rich to sulfate-rich after probable acidification by sulfur oxidation.
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correspond to the Ca-rich path. Almost all high thermal inflow waters (Twater Tatm>20 jC) belong to the SO4-rich group. Only 3 out of 22 hot inflow waters follow the Ca-rich path. It must be emphasized that the evolutionary path is the only parameter for which some relation with temperature has been detected so far. None of the dissolved components presents any relation with temperature (see Section 7.7). 9.3.4. Discrepancy between lithology and brine evolution The geographical distribution of SO4-rich and Carich salars is shown in Fig. 19. According to the previous discussion, salars located in volcanic basins are expected to contain SO4-rich brines, while those surrounded by sedimentary rocks should be filled with Ca-rich brines, as perfectly illustrated in salar de Atacama [13]. In addition to Atacama, two other salars are located in the Pre-Andean Depression. where sedimentary formations of the Precordillera predominate: Imilac [29] and Punta Negra [30]. Both belong to the Ca-rich type. All other salars, except Pintados [5], are in volcanic basins of the Western Cordillera and are expected to be of the SO4-rich type as a result of the abundance of native sulfur in most of the volcanoes. This is effectively observed in the north of the study area (18 – 21jS latitude). Salar de Pintados [5], in the Central Valley, is mainly fed by inflows originating in the volcanic Western Cordillera. In the central area (21 –24jS latitude), both types are intermingled. No simple geological argument can explain the difference from one basin to the other. In the south of the study area (25 – 27jS latitude), salars are predominently of the Ca-rich type close to the Precordillera and of the SO4-rich type eastward in the volcanic Cordillera. The chemistry of Ca-rich salars located in the volcanic Western Cordillera seems to be controlled by sedimentary formations. If some of them are adjacent to the Precordillera (Pedernales [44], Maricunga [52]), others are well inside the Western Fig. 19. Outline map of north Chilean Andes showing the chemistry of each salar: SO4-rich, Ca-rich or both. In the south, Ca-rich salars predominate close to the Precordillera, while SO4-rich salars are found eastward within the Cordillera (compare with Fig. 12). Note the slightly exaggerated horizontal scale.
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Cordillera, up to 50 km from its western boundary (Aguas Calientes 2 [21]; Wheelwright [49]). It is difficult to conceive a major contribution of sedimentary rocks, either from the west or from depth, in salars located in such massive volcanic formations. A much better explanation is the recycling of present infiltrating brines from the salars themselves (see Section 7). However, such an explanation only pushes the problem further back in time. When did the waters acquire their Ca-rich composition? The drainage basins of the Ca-rich salars, now mostly or exclusively volcanic, must have been sedimentary in the past, before the establishment of recent Pleistocene volcanoes and ignimbrite sheets. Consequently, these salars should be notably old, of Pleistocene age or even earlier. Stoertz and Ericksen (1974) have already
observed that many salars in the Western Cordillera have been tilted downward to the northwest during the Quaternary. Salares de Pedernales [44] and Maricunga [52] are precisely among the most asymmetric ones, which points to an older age than that of nontilted salars within the Western Cordillera. The ancient Carich underground waters may have been continuously recycled up to now. The alteration of volcanic rocks progressively shifts the chemical type of the solutions. Both Pedernales and Maricunga belong to the mixed Ca/SO4 salar group. Inflow waters result from the mixing in variable proportion of recycled brines with dilute weathering solutions. The weathering of sulfur-rich volcanic rocks increases the SO4 content, and then the SO4/Ca ratio, of inflow waters. Alternatively, if the
Fig. 20. Very schematic representation of the recycling and mixing processes responsible for the chemical shift from Ca-rich to SO4-rich brines in Chilean salars. For the sake of simplicity, we have restricted the cycle to one basin. Actually, deep water circulation may transport infiltrating salts in several distant basins. Size of symbols only reflects the proportion of Ca and SO4 in each solution. The salar brine is of the Ca-rich type, which means that the concentration of Ca is markedly higher than that of sulfate (as suggested by the thick bold characters used for Ca). The infiltrated brine has the same composition. Heat flow recycles this Ca-rich brine along the Ghyben – Herzberg mixing zone, in contact with fresh waters that have leached volcanic rocks of the drainage basin. Sulfur-rich volcanoes produce SO4-rich waters which mix with Ca-rich recycled brines. Thus, the Ca/SO4 ratio of the mixed solution is lower than that of the salar brine (spring 1). The gap between Ca and SO4 concentration is reduced. The weathering of sulfur-poor volcanoes produces alkaline waters. Calcite is a common weathering product whose precipitation lowers the Ca concentration of the water. Here again, the mixing of Ca-poor waters with Ca-rich infiltrated brines reduces the gap between Ca and SO4 concentration in the resulting inflow water (spring 2). Therefore, on the one hand, inflows tend to reduce the excess of Ca over SO4 in the salar brine; on the other hand, the evaporative concentration amplifies the difference between Ca and SO4 in the brine. A long time is likely to be necessary to shift the Ca-rich brine composition to a SO4-rich composition.
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sulfur content of volcanic rocks is low, then the weathering solutions are enriched in HCO3 (see Section 9.1), which also reduces the Ca concentration of the mixed inflow waters due to calcite precipitation. Here again, the SO4/Ca ratio of the inflow water increases. However, the evaporative concentration in the basin amplifies the difference on behalf of the most concentrated component (Ca). Thus, the SO4/Ca ratio of the infiltrating brine is much lower (more Ca) than that of the input water. This amplification mechanism may delay for a long time the shift from the Carich to the SO4-rich type. Eventually, SO4 overtakes Ca in spring waters and the amplification process then favors the rapid increase of SO4 in lake brines. Fig. 20 summarizes the whole cycle.
10. Concluding remarks Three main sources of salts have been detected in inflow waters of Chilean salars. Atmospheric inputs, through precipitation and dry fallout, affect the whole area but are noticeable only in a few salars fed by very dilute inflows [below 100 mg/l total dissolved solids (TDS)]. Most of the minor components in very dilute inflows, and a significant part of the major ones, stem from sea salts and desert dust, especially NO3, Br and As. Volcanic rock alteration is the main source of salts in inflow waters with concentration ranging roughly from 100 to 600 mg/l TDS. However, the average salt content of all inflows (3230 mg/l) is much higher than that derived from the mere alteration of volcanic rocks. Brine recycling is the main source of salts in most salars. Saline lakes are under steady-state regime. The residence time of conservative components ranges from a few years to some thousands years. Most of the dissolved salts brought in the lakes by inflow waters are lost by infiltration through bottom sediments and return to the hydrologic cycle. The infiltrating brines provide the major part of dissolved salts to underground waters. The dissolution of surficial salt crusts also contributes to the salt load of deep waters. Owing to the complexity of water circulation in fractured volcanic formations, infiltrating brines may be transported to adjacent basins, or even further. Thermal convection due to the high heat flow is the likely driving force that moves the waters. Thermal heating affects about 78% of springs and
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underground waters. Thermal heating may have different results below the salar and below the mountains. Below the salar, heating could tend to oppose the downward brine flux and favor recycling in the same salar. In contrast, below the volcanic formations, the heating could increase the loss of water from the basin to other basins (see Fig. 20). Ancient evaporites, especially halite and gypsum, are present in depth below the volcanic formations. However, the Cl/Br ratio of halite is much higher than that of inflow waters, which precludes the dissolution of deep salt rock in underground waters. On the contrary, the Cl/ Br ratio confirms the mixing of present lake brines with snow meltwaters as the origin of saline inflows. Only ancient gypsum seems to contribute some Ca and SO4 to inflow waters. Stable isotopes also supports the mixing hypothesis: the d2H and d18O values of inflow waters plot along a straight line, probably a mixing line, joining snow meltwaters to lake brines. Chilean salars in the Andean Cordillera cannot be strictly considered as closed basins or terminal lakes. Most of the salts entering the salars are lost by infiltration. Only poorly soluble salts, such as calcite, Mg salts and gypsum, remove dissolved components which slowly accumulate in saline sediments. The high concentration of many inflows does not reflect an intense material transfer by weathering, evaporite dissolution or hydrothermal alteration. Nor does it mean a high accumulation rate of salts in the basins. Actually, true material transfer is mostly restricted to atmospheric inputs and volcanic rock weathering. This is intuitively understandable: if all inflowing salts would accumulate in Andean salars, very thick salt crusts should be observed. Thick modern salt crusts are not found in this region. According to the model of brine evolution observed in most saline lakes worldwide, three main groups of brines may be generated by evaporation and chemical divide (Hardie and Eugster, 1970): alkaline brines (Na/HCO3 – CO3 – Cl, pH>9, traces of Ca and Mg), sulfate-rich brines (Na/SO4 –Cl, low Ca, pH < 9) and Ca-rich brines (Na – Ca/Cl, low SO4, pH < 9). In northern Chile, alkaline salars are almost completely lacking. Only one brackish pond out of 227 was found to be alkaline. All others belong to the SO4-rich and Ca-rich groups. This paucity is due to the abundance of native sulfur in the Western Cordillera and to atmospheric deposition of gypsum-rich desert dust
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from the Central Valley. The oxidation of sulfur produces SO4 and acidifies the inflow waters, reducing drastically their carbonate content. Gypsum dust enriches the waters in Ca, which conversely lowers their carbonate contents because of the very low solubility of calcite. Therefore, the (SO4 + Alk/2)/Ca ratio was found to be a valuable parameter to class and investigate Chilean salars. The abundance of native sulfur in volcanoes of the Western Cordillera leads to sulfate-rich inflow waters, which in turn produce sulfate-rich brines by evaporative concentration. Conversely, sedimentary rocks of the Precordillera contain Ca minerals (calcite, dolomite, gypsum) whose weathering leads preferentially to Ca-rich waters and Ca-rich brines. Small differences of SO4 and Ca concentration in dilute inflows are amplified in the final brines by evaporative concentration. This overall relation between lithology and brine composition is remarkably illustrated in the largest Chilean salar: the salar de Atacama. However, several salars in volcanic basins of the Western Cordillera belong to the Ca-rich group. Their chemistry may have been inherited from the weathering of ancient sedimentary formations during the late Tertiary or the Pleistocene. Later on, ignimbrite sheets and lava flows may have covered the sedimentary formations. Such a process may well account for Carich salars adjacent to the Precordillera, at the western fringe of the volcanic cordillera. Most of the salt load of inflow waters does not stem directly from rock alteration in the drainage basins, but from lake brine recycling. The recycling process is likely to be as old as the salars themselves. Infiltration rates tend to slowly decrease with time because of the progressive choking of bottom sediments by clay and silt particles. Thus, brine leakage was probably more effective in past times. Nonequilibrium conditions seem to control important features of Chilean salars. Old salt crusts deposited under very arid climate are presently undergoing dissolution in the semiarid Cordillera. The Ca-rich composition of many waters and brines does not reflect the volcanic lithology of the Western Cordillera. The global hydrogeochemical cycle is not yet equilibrated with the presently dominant volcanic lithology. The quantity of weathering material entering the hydrological cycle is much lower than the quantity of recycled salts.
A complex correlation between lithology, thermalism and salar chemistry is hinted at in the south of the study area. Ca-rich salars are preferentially located close to the sedimentary Precordillera and their inflow waters undergo little thermal influence. Conversely, SO4-rich salars tend to be located well inside the volcanic Western Cordillera and are largely fed by thermal waters. This is consistent with the sedimentary to volcanic transition from west to east associated to the increase of heat flow in the same direction. Obviously, many factors obscure this overall correlation: the sedimentary formations contain volcanic and volcano-sedimentary units. The boundary between Precordillera and Western Cordillera is not so precisely delimited as shown in our schematic maps. Thermal waters may be cooled by mixing with fresh meteoric waters. Deep circulation in fractured eruptive rocks can be very complex. The northern and central areas seem to be particularly affected by these drawbacks. Dilute meteoric waters are more contaminated by salt recycling in areas containing many salars. Salts lost from one salar may be widely redistributed in several other basins. As a result, exploration for fresh water should be focused in areas devoid of salars. In salar basins, the water table is close to the topographic surface, which reduces significantly the drilling costs. The drawback is that such waters are often of poor quality. Outside salar areas, waters are much deeper, but probably less affected by salt contamination.
Acknowledgements We want to express our deep gratitude to the institutions and companies that gave us invaluable and unselfish logistic assistance: the regional Bureaus of the Direccio´n General de Aguas (Regiones I, II, III); Eje´rcito de Chile; Minera Escondida; Codelco Chile: Division Chuquicamata and Division Salvador; Minera Mantos Blancos; Corporacion Nacional Forestal; Minera Quiborax; Minera Don˜a Ine´s de Collahuasi; and Minera Zaldivar. Thanks are due to Professor Abraham Lerman and to an anonymous reviewer who made constructive reviews of the manuscript. This is EOST contribution no. 2003.403-UMR7517.
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Appendix A (continued)
Appendix A Cl (mmol/l) and Br (Amol/l) of waters and brines of Chilean salars. Other components in Risacher et al. (1999). N = snow, S = spring, P = seep, R = river, L = lake, U = underground water.
Name
Type
CL/BR
NEV-1 NEV-2 NEV-3 NEV-4 NEV-5 NEV-6 NEV-8 NEV-9 AC1-1 AC1-4 AC1-6 AC1-8 AC1-11 AC2-1 AC2-4 AC2-7 AC2-10 AC2-13 AC2-15 AC3-1 AC3-4 AC3-7 AC3-10 AC3-12 AC3-13 AC3-14 AC3-15 AC4-1 AC4-2 AC4-3 AC4-4 AC4-5 AC4-6 AC4-7 AC4-8 AC4-9 AC4-10 AGI-1 AGI-1A AGI-2 ALC-1 ALC-2 ALC-3 ALC-4 ALC-5 ALC-6
N N N N N N N N S S L L S S P S P P L P L S P P P S L P L L P L L L P S S L L L U S S S S L
0.015/0.05 0.021/0.09 0.018/0.13 0.029/0.10 0.025/0.08 0.003/0.02 0.003/0.04 0.005/0.02 408/56 102/18 1820/280 2020/210 17.3/7.3 19.4/2.6 35.5/7.7 40.0/12 217/28 35.2/13 165/45 75.0/14 361/46 28.0/11 126/13 22.3/4.2 116/40 34.1/14 48.0/18 9.39/5.0 380/190 88.3/43 27.1/12 271/130 892/460 4930/2400 25.2/16 5.22/2.8 16.9/10 2960/230 3200/230 5880/150 2.52/2.7 0.169/0.49 0.168/0.57 0.197/0.30 0.184/0.40 1030/370
Name
Type
CL/BR
ALC-7 ALC-8 ALC-9 ALC-10 AMA-1 AMA-2 AMA-3 AMA-4 AMA-5 AMA-6 ASC-1 ASC-2 ASC-3 ASC-4 ASC-7 ASC-9 ASC-10 ASC-13 ASC-14 ASC-15 ASC-16 ASC-17 ASC-18 ASC-19 ASC-20 ASC-21 ASC-22 ASC-23 ASC-24 ASC-25 ASC-26 ASC-27 ASC-28 ASC-29 ASC-30 ASC-31 ASC-32 ASC-33 ASC-34 ASC-36 ASC-37 ASC-38 ASC-39 ASC-40 ASC-41 ASC-42 ASC-43 ASC-44 ASC-45 ASC-46 ASC-47 ATA-1 ATA-2
P L L S L L P L S L U U U U U U U S S U U U L S L L U L U L S L S L S L L S L S S S L L L P L S L L U R S
2.03/2.7 14.9/24 58.6/110 0.328/1.0 538/31 284/17 129/9.2 3410/200 279/52 1420/68 32.7/18 8.88/1.5 25.3/15 30.5/9.4 109/8.6 35.5/9.0 60.1/10 0.087/0.03 0.081/0.05 25.0/14 3.01/1.9 47.2/13 211/43 122/7.9 124/9.0 287/21 31.6/7.7 91.3/19 31.4/8.7 67.5/21 30.9/6.0 64.7/15 56.3/8.7 151/23 54.4/10 136/21 657/150 55.8/9.7 199/41 49.1/18 33.9/17 27.8/13 266/120 107/24 1650/550 28.4/1.7 916/170 3.69/0.75 583/120 245/68 30.3/10 11.8/0.48 2.32/0.28 (continued on next page)
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Appendix A (continued)
Appendix A (continued)
Name
Type
CL/BR
Name
Type
CL/BR
ATA-3 ATA-4 ATA-5 ATA-6 ATA-7 ATA-8 ATA-9 ATA-10 ATA-11 ATA-12 ATA-13 ATA-14 ATA-15 ATA-16 ATA-17 ATA-18 ATA-19 ATA-20 ATA-21 ATA-22 ATA-23 ATA-24 AZU-1 AZU-2 AZU-3 AZU-4 AZU-5 AZU-6 AZU-7 BAY-1 BAY-2 BRA-1 BRA-2 BRA-3 BRA-4 CAR-1 CAR-2 CAR-3 CAR-4 CAR-5 CAR-6 CAR-7 CAR-8 CAR-9 CAR-10 CAR-11 CAR-12 CAR-13 CAR-15 CAR-16 CAR-17 CAR-18 CAR-19 CAR-20
R R U R R R R R R R R R R S R R P P P U S S P L L L L L U L P P L P L S S L L S L L L L L L S L S L L S L L
7.10/1.6 17.0/1.3 27.0/18 1.50/1.0 8.50/5.4 2.03/5.0 1.25/1.0 6.50/3.8 2.50/1.6 25.0/11 43.0/11 15.4/10 15.0/10 15.0/10 3.50/2.1 0.480/0.42 10.5/7.4 304/40 25.3/11 52.9/14 7.76/4.2 0.041/0.08 27.4/6.4 893/78 2190/280 120/23 292/56 5700/870 4.41/1.1 28.9/23 15.9/7.1 35.9/15 1810/790 21.8/11 47.7/24 0.269/0.37 0.272/0.32 1030/59 3020/130 47.1/3.1 2020/95 5700/530 2900/190 31.7/2.3 190/8.8 1250/70 126/13 1800/190 17.4/2.3 5490/130 463/87 11.4/2.6 273/48 36.8/7.6
CAR-21 CAR-22 CAR-24 CAR-25 CAR-26 CHR-1 CHR-2 CHR-3 CHR-4 CHR-5 CHR-6 CHR-7 CHR-8 COP-1 COP-2 COP-3 COP-4 COP-5 COP-6 COP-7 COP-8 COP-9 COP-10 COP-11 COP-13 COP-14 COP-15 COP-16 COP-17 COP-18 COP-19 COP-20 COP-21 COP-22 COP-23 COP-24 COT-1 COT-2 COT-3 COT-4 CPR-1 CPR-2 CPR-3 CPR-4 CPR-5 CPR-6 CPR-7 ESC-1 ESC-2 ESC-3 ESC-4 ESC-5 ESC-6 FRA-1
L S P S L L R S S L S L S S S U P S L U U U L L L L L U S L L L L U U U S L L L L S L S L L L L L P P L L R
90.5/11 6.74/2.2 109/8.4 419/22 1250/69 2.71/0.57 0.091/0.09 0.007/0.05 0.045/0.06 0.087/0.08 0.046/0.05 1.74/0.23 0.016/0.02 1.72/1.8 43.9/25 0.269/0.38 5.30/2.0 4.89/2.2 461/180 28.5/4.3 0.777/0.78 21.2/2.6 461/160 1430/29 32.3/12 121/42 5.60/2.4 20.9/2.2 13.7/3.6 807/190 187/46 43.6/10 4410/830 3.01/1.6 1.13/1.0 0.739/0.89 0.168/0.14 0.838/0.32 0.903/0.16 2.55/1.3 197/25 106/17 334/40 102/16 847/76 1190/100 3770/280 41.1/33 199/180 18.7/14 30.5/30 347/330 71.6/59 0.483/0.17
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293 Appendix A (continued)
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Appendix A (continued)
Name
Type
CL/BR
Name
Type
CL/BR
FRA-2 FRA-3 FRA-4 FRA-5 FRA-6 FRA-7 FRA-8 FRA-9 FRA-10 FRA-12 FRA-13 FRA-14 FRA-15 FRA-16 FRA-17 FRA-18 GOR-1 GOR-2 GOR-3 GOR-4 GOR-5 GOR-6 GOR-7 GOR-8 GOR-9 GOR-10 GOR-11 GOR-12 GRA-1 GRA-2 GRA-3 GRA-4 HCO-1 HCO-2 HCO-3 HCO-4 HCO-5 HCO-6 HCO-7 HCO-8 HCO-9 HCO-10 HCO-11 HCO-12 HCO-13 HCO-14 HCO-15 HCO-16 HCO-17 HCO-18 HEL-1 HEL-3 HEL-5 HEL-6
P R L P S L S L S P S S L R R S R P L L L P P L L P L L P L L L S R R R R U S R U R S S L L L L L L S L S S
2.73/1.3 0.023/0.03 4810/2700 0.608/0.41 0.405/0.07 429/170 0.630/0.10 97.0/37 3.41/2.0 11.4/7.5 1.02/0.49 0.339/0.30 20.1/3.2 3.05/2.5 0.322/0.14 83.1/88 3.36/0.72 54.5/41 111/86 258/48 1350/230 5.63/2.0 112/8.6 104/100 506/270 46.4/52 3450/110 3490/80 119/12 305/30 2120/170 440/47 0.819/1.2 0.145/0.25 0.269/0.64 0.256/0.47 0.748/0.72 2.30/2.6 0.100/0.40 0.088/0.36 0.576/0.73 0.390/0.64 1.43/0.98 0.916/1.3 140/78 222/140 80.2/47 12.5/8.0 246/150 909/470 1.66/0.24 5050/940 0.250/0.31 0.109/0.18
HEL-7 HEL-8 HEL-9 IGN-1 IGN-3 IGN-4 IGN-5 IMI-1 IMI-2 IMI-3 IMI-5 IMI-6 IMI-7 IMI-8 IMI-8D INF-1 INF-2 INF-3 INF-4 INF-5 INF-6 ISL-1 ISL-2 ISL-3 ISL-4 ISL-5 ISL-6 ISL-7 ISL-8 ISL-9 ISL-10 ISL-11 ISL-12 ISL-13 ISL-14 ISL-15 ISL-16 ISL-17 ISL-18 ISL-19 JIL-1 JIL-2 JIL-3 JIL-4 JIL-5 JIL-6 LAA-1 LAA-2 LAA-3 LAA-4 LAA-5 LAA-6 LAA-7
S L L P L L L U L U L L L P U P L L L L P S L L P L L L U S L L L S L L S L S L L P L L L P U R R R S R S
0.537/0.35 23.7/5.6 14.9/22 4.00/1.1 17.0/1.8 12.5/1.9 39.6/5.1 2.70/2.0 3.24/2.1 55.5/7.8 1410/8.4 713/8.9 507/11 150/5.7 181/6.9 15.3/6.2 142/41 5540/250 1230/97 280/28 119/19 276/2.2 5310/28 1410/8.0 102/1.1 1540/11 425/2.2 818/11 155/3.6 262/4.3 658/11 2130/21 5070/22 494/2.5 2920/16 1220/8.0 226/2.5 5030/38 88.3/1.5 2470/52 177/100 9.93/7.3 115/67 123/67 215/110 32.3/13 0.101/0.29 1.14/1.1 0.964/0.85 1.52/1.4 0.682/0.45 0.956/0.51 0.246/0.32 (continued on next page)
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Appendix A (continued)
Appendix A (continued)
Name
Type
CL/BR
Name
Type
CL/BR
LAA-8 LAC-1 LAC-4 LAC-5 LAC-7 LAV-1 LAV-2 LAV-3 LAV-4 LAV-5 LAV-6 LAV-7 LAV-8 LAZ-1 LAZ-2 LAZ-3 LAZ-4 LEJ-1 LEJ-3 LGN-1 LGN-2 LGN-3 LGN-4 LGN-5 LGU-1 LGU-2 LGU-3 LOY-1 LOY-5 LOY-9 LOY-13 LOY-16 LOY-20 LOY-24 LOY-28 LOY-32 LOY-35 LOY-39 LOY-40 LOY-41 MAR-1 MAR-2 MAR-3 MAR-4 MAR-5 MAR-6 MAR-7 MAR-8 MAR-9 MAR-10 MAR-11 MAR-12 MAR-13 MAR-14
R P U L P L R P L L S S R S L L P L P P L P L S S L U R P P R P L P P U P L S S S R R R R U R R R L L P P L
0.268/0.38 5.43/1.8 5.22/1.5 325/45 24.2/5.9 2580/1300 7.60/3.8 78.7/72 56.4/58 520/220 41.1/16 8.02/3.1 24.0/12 37.5/28 1580/990 166/110 28.0/15 274/130 3.14/1.0 22.3/9.3 73.9/41 100/54 232/160 28.6/30 0.906/1.2 4.53/3.7 0.268/0.88 38.4/3.5 230/19 1.50/0.53 0.515/0.33 3.38/0.68 64.3/6.2 108/20 59.0/4.8 57.2/10 62.0/5.5 4170/65 1.11/0.51 0.489/0.26 0.501/0.22 0.856/0.63 5.29/3.1 29.5/17 1.00/0.07 2.44/2.4 12.2/7.3 0.096/0.04 2.20/1.2 61.0/14 3.48/2.0 35.3/19 19.1/9.7 251/32
MAR-15 MAR-16 MAR-17 MAR-18 MAR-20 MAR-21 MAR-22 MAR-23 MAR-24 MAR-25 MAR-26 MAR-27 MAR-28 MAR-29 MAR-30 MIC-1 MIC-2 MIC-3 MIC-4 MIC-5 MIC-6 MIC-7 MIC-8 MIC-10 MIC-13 MIC-14 MIC-15 MIN-1 MIN-3 MIN-6 MIN-8 MIS-1 MIS-3 MIS-5 MIS-8 MIS-9 MUE-1 MUE-4 PAJ-1 PAJ-2 PAJ-3 PAJ-4 PAJ-5 PAJ-6 PAJ-7 PAJ-8 PAJ-9 PAJ-10 PAJ-11 PAR-1 PAR-2 PAR-3 PAR-4 PAR-5
P U U U L P L L L L P S U P R S R S L L L L U U U S L L L P L P L P L U P L S L U L L L P S L L L P L L P L
13.8/2.3 62.1/25 26.1/15 0.838/0.54 130/13 38.5/6.3 66.8/12 159/38 1420/280 5750/1300 22.7/3.9 1.12/0.85 1.20/0.27 0.185/0.05 10.6/4.0 0.138/0.23 0.594/0.48 0.664/0.60 3.95/2.4 181/98 20.9/8.8 50.0/22 0.511/0.57 0.661/0.75 0.397/0.65 0.157/0.44 5.21/5.5 74.6/30 67.1/24 23.4/9.0 62.6/24 1.57/0.90 32.4/13 3.86/1.4 34.0/12 0.089/0.37 364/38 1090/130 240/38 907/140 371/53 723/120 4250/620 2820/270 381/35 193/59 1890/670 3720/820 2390/690 139/3.6 568/15 3710/87 254/6.8 5290/180
F. Risacher et al. / Earth-Science Reviews 63 (2003) 249–293 Appendix A (continued)
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Appendix A (continued)
Name
Type
CL/BR
Name
Type
CL/BR
PAR-6 PAR-7 PED-1 PED-2 PED-3 PED-4 PED-5 PED-6 PED-7 PED-9 PED-10 PED-11 PED-12 PED-13 PED-14 PED-15 PED-16 PED-17 PED-18 PED-19 PED-20 PED-21 PED-22A PED-22 PED-23 PED-25 PED-26 PED-27 PED-28 PED-29 PIE-1 PIE-2 PIE-3 PIE-4 PIE-5 PIE-6 PIE-7 PIE-8 PIE-9 PIE-10 PIE-11 PIE-12 PIE-13 PIE-14 PIN-1 PIN-2 PIN-3 PIN-4 PIN-5 PIN-6 PIN-8 PIN-9 PIN-10 PIN-11
S L S R P R S S U U P L P S L L P S S L U U L L L U U U U U P L L P L L L P L L S L P L U U U S U U S R S U
133/4.3 708/23 28.2/22 29.9/13 2.87/1.1 8.62/2.5 6.43/0.47 61.0/6.1 1.81/0.93 0.070/0.05 2.89/1.6 3640/190 1310/68 40.3/2.4 1790/50 5450/97 245/8.7 4.87/5.6 0.799/0.67 5410/180 0.157/0.16 0.148/0.14 175/35 510/97 1010/120 0.252/0.27 35.2/12 33.3/12 40.9/16 35.5/13 41.6/8.2 353/75 97.5/22 84.4/29 304/110 492/130 269/49 38.8/6.9 1220/190 163/38 30.0/6.9 419/170 38.8/7.3 141/26 2.10/2.3 3.70/3.1 2.60/2.7 6.00/5.0 5.96/5.3 17.4/17 5.26/4.4 74.5/45 1.80/2.1 23.4/18
PIN-12 PIN-13 PIN-14 PIN-15 PIN-16 PIN-17 PIN-18 PIN-19 PIN-20 PIN-21 PIN-22 PIN-23 PIN-24 PIN-25 PIN-26 PIN-27 PIN-28 PIN-29 PIN-30 PIN-31 PIN-32 PIN-33 PIN-34 PIN-35 PIN-36 PIN-37 PIN-38 PIN-39 PIN-40 PIN-41 PIN-42 PIN-43 PIN-44 PIN-45 PSA-1 PSA-5 PSA-8 PSA-11 PUN-1 PUN-2 PUN-3 PUN-4 PUN-5 PUN-6 PUN-7 PUN-8 PUN-9 PUN-10 PUN-11 PUN-12 PUN-13 PUN-14 PUN-15
S U S U U U S S S S S S S U U U U U R R S S R U L R S S U L U L L L R L P P L L L P L L L U L P U S S S U
0.740/1.4 0.931/1.6 2.49/2.4 3.12/3.6 19.4/16 4.09/4.5 1.22/2.3 1.47/2.7 1.18/0.44 1.15/2.2 3.01/4.5 1.14/2.2 3.98/6.5 6.82/6.7 19.1/17 21.9/19 23.6/21 6.97/6.5 32.3/29 1.06/2.1 1.88/2.9 1.51/2.0 7.69/5.8 10.0/7.9 935/120 3.65/3.4 0.391/1.8 0.406/1.7 7.08/6.6 236/39 12.3/9.3 4390/680 1440/190 115/24 0.983/0.61 343/230 6.36/3.5 0.559/0.20 309/92 195/57 162/47 170/11 294/21 4010/120 155/3.8 112/2.2 706/17 238/9.8 33.6/11 2.12/2.4 5.17/2.8 35.6/31 37.9/32 (continued on next page)
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Appendix A (continued)
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Name
Type
CL/BR
PUN-16 PUN-17 PUN-19 PUN-21 PUN-22 PUN-23 SUR-1 SUR-2 SUR-3 SUR-4 SUR-5 SUR-7 SUR-8 SUR-9 SUR-10 SUR-11 SUR-12 SUR-13 SUR-14 SUR-15 SUR-16 SUR-17 SUR-18 SUR-19 SUR-20 SUR-21 SUR-22 SUR-23 SUR-24 SUR-25 SUR-26 TAR-1 TAR-3 TAR-5 TAR-6 TAR-9 TAR-11 TAR-12 TAR-13 TRI-1 TUY-1 TUY-4 WHE-1 WHE-2 WHE-3 WHE-4 WHE-5 WHE-6 WHE-7 WHE-8 WHE-9
U R U U S R R S P S P S L S P P S L S P S L S L S L L L L L S S L R P L S S R L R L R L S L L L P L P
43.9/36 43.4/32 4.84/0.77 6.48/4.8 0.184/0.28 0.491/0.35 0.141/0.19 47.9/36 3.27/2.9 1.49/1.2 6.52/4.0 0.653/0.97 19.0/14 39.3/32 23.2/19 0.793/0.79 1.20/1.0 1040/680 0.514/0.85 0.526/0.57 6.12/4.6 614/380 0.375/0.53 459/300 226/170 967/680 50.6/33 4380/2700 1970/1300 89.3/66 0.286/0.59 8.60/4.0 34.8/15 1.60/1.1 7.25/3.5 2760/1100 9.18/5.0 26.4/13 15.7/6.8 131/14 88.9/9.7 2120/160 100/11 2750/270 48.7/5.2 329/42 2480/290 708/96 96.1/12 175/25 20.2/3.2
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aquifer, Rainer Mesa, Nevada. U.S. Geological Survey WaterSupply Paper 1535-Q. Yechieli, Y., Wood, W.W., 2002. Hydrogeologic processes in saline systems: playas, sabkhas, and saline lakes. Earth-Science Reviews 58, 343 – 365. Yonts, W.L., Giese, G.L., Hubbard, E.F., 1973. Evaporation from Lake Michie, North Carolina, 1961 – 71. U.S. Geological Survey Water-Resources Investigations 38-73, 27 pp. Francß ois Risacher is a graduate of the Ecole Nationale Supe´rieure de Ge´ologie de Nancy (1971). He holds a PhD in Geochemistry from the University Louis Pasteur of Strasbourg (1992). He is a Senior Researcher at the Institut de Recherche pour le De´ veloppement (a French cooperation agency formerly known as Orstom). For more than 20 years, he carried out applied and fundamental research on evaporitic basins in Bolivia and Chile, focusing his interest on economical resources and geochemistry of saline lakes and salt crusts. Hugo Alonso is a Chemical Engineer from the Universidad Cato´lica del Norte (UCN), Chile (1965). Afterwards, he studied Soils and Hydrological Sciences in France. His research interests focus on soil reclamation in the Atacama Desert, irrigation with saline waters in native Andean communities, arsenic and boron contamination in water and crops, preservation of water quality in mining areas. In 1995, he received a National Government Award for his contributions to the use of waters in the arid zone of northern Chile. Now, he is Professor of Environmental Geochemistry at UCN and a consultant on water quality management for the National Environmental Commission, a Chilean government agency, and for several mining companies. Carlos Salazar is a Civil Engineer from the Universidad de Chile (1981). He joined the Chilean National Water Resources Service (Direccion General de Aguas), where he is principally involved in water resources development and global management as the Head of the Studies and Planning Department. He has special interests in groundwaters and wetland systems in arid areas.