Soil carbon distribution and site characteristics in hyper-arid soils of the Atacama Desert: A site with Mars-like soils

Available online at www.sciencedirect.com

Advances in Space Research 50 (2012) 108–122 www.elsevier.com/locate/asr

Soil carbon distribution and site characteristics in hyper-arid soils of the Atacama Desert: A site with Mars-like soils Julio E. Valdivia-Silva a,b,⇑, Rafael Navarro-Gonza´lez a, Lauren Fletcher c, Sau´l Perez-Montan˜o d, Renee´ Condori-Apaza e, Christopher P. Mckay b a

Laboratorio de Quı´mica de Plasmas y Estudios Planetarios, Instituto de Ciencias Nucleares, Universidad Nacional Auto´noma de Me´xico, Circuito exterior S/N, Ciudad Universitaria, Apartado Postal 04510, Me´xico DF, Mexico b Space Sciences Division, NASA Ames Research Center, Moffett Field, CA 94035, USA c University of Oxford, Atmospheric, Oceanic and Planetary Physics, Oxford, UK d Department of Chemistry, San Jose State University, San Jose, CA 95192, USA e Universidad Nacional San Agustı´n, Facultad de Ingenierı´a Quı´mica, Arequipa, Peru Received 25 September 2011; received in revised form 1 March 2012; accepted 2 March 2012 Available online 13 March 2012

Abstract The soil carbon content and its relation to site characteristics are important in evaluating current local, regional, and global soil C storage and projecting future variations in response to climate change. In this study we analyzed the concentration of organic and inorganic carbon and their relationship with in situ climatic and geological characteristics in 485 samples of surface soil and 17 pits from the hyper-arid area and 51 samples with 2 pits from the arid–semiarid region from the Atacama Desert located in Peru and Chile. The soil organic carbon (SOC) in hyperarid soils ranged from 1.8 to 50.9 lg C per g of soil for the 0–0.1 m profile and from 1.8 to 125.2 lg C per g of soil for the 0–1 m profile. The analysis of climatic (temperature and precipitation), elevation, and some geologic characteristics (landforms) associated with hyper-arid soils explained partially the SOC variability. On the other hand, soil inorganic carbon (SIC) contents, in the form of carbonates, ranged from 200 to 1500 lg C per g of soil for the 0–0.1 m profile and from 200 to 3000 lg C per g of soil for the 0–1.0 m profile in the driest area. The largest accumulations of organic and inorganic carbon were found near to arid–semiarid areas. In addition, the elemental carbon concentrations show that the presence of other forms of inorganic carbon (e.g. graphite, etc.) was negligible in these hyperarid soils. Overall, the top 1 m soil layer of hyperarid lands contains 11.6 Tg of organic carbon and 344.6 Tg of carbonate carbon. The total stored carbon was 30.8-fold the organic carbon alone. To our knowledge, this is the first study evaluating the total budget carbon on the surface and shallow subsurface on 160,000 km2 of hyperarid soils. Ó 2012 COSPAR. Published by Elsevier Ltd. All rights reserved. Keywords: Carbon storage; Hyperarid soils; Atacama Desert; Pampas de La Joya; Mars analogue

1. Introduction The Atacama Desert is located along the western coast of South America throughout the shore region of southern Peru and northern Chile covering about 3500 km, between ⇑ Corresponding author. Address: NASA Ames Research Center, Mail Stop: 245:3, Office 213A, Moffett Field, CA 94035, USA. Tel.: +1 650 604 1136; fax: +1 650 604 6779. E-mail address: [email protected] (J.E. Valdivia-Silva).

10°S to 35°S latitude, and 70°W to 72°W longitude. Because the average values of precipitation in the complete region are less than 200 mm/y, Houston and Hartley (2003) divided this desert according to the aridity index (AI) as semiarid (0.2 < P/PET < 0.5), arid (0.05 < P/PET < 0.2), and hyper-arid (P/PET < 0.05) regions (Fig. 1, Table 1, & Table 1S). This index was calculated as the ratio of precipitation and potential evapotranspiration (P/PET) using Thornthwaite’s equations as a function of mean monthly temperatures and mean monthly number of daylight hours (Thornwaite,

0273-1177/$36.00 Ó 2012 COSPAR. Published by Elsevier Ltd. All rights reserved. doi:10.1016/j.asr.2012.03.003

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

Fig. 1. Location of the Atacama Desert in western South America showing dominant climatic regions. The limits of each region were corroborated using field data (see text) and published studies by Houston and Hartley (2003). Boxes show the most important dry cores in the peruvian and chilean parts from the desert (Fig. 2A and B).

1948; UNEP, 1997). Hyper-arid areas emerge independent of human activities under a natural development and evolution of drier climates denominated “aridization”. This process takes place much more slowly than the processes of “desertification” (Kottek et al., 2006), which are defined as the land degradation in dry areas resulting from climatic variations and human activities (Houerou, 1996). Thus, hyper-arid lands are usually excluded from the consideration of desertification (Schlesinger, 1997). The combined effects of a high pressure system located on the western Pacific Ocean, the cold north-flowing Humboldt Current, and the rain shadow of the Andean Cordillera intercepting precipitation from the inter-tropical convergence are the main factors involved in the hyper-arid climate formation in this region (Arroyo et al., 1988; Houston and Hartley, 2003). These factors have been used in support of geological evidence that the Atacama Desert has remained hyper-arid between 9 and 25 Ma (Alpers and Brimhall, 1988; Dunai et al., 2005; Evenstar et al., 2005). Studies based on sedimentologic evidence estimate that the beginning of continuous hyper-aridity occurred until the Late Pliocene (Hartley and Chong, 2002; Hartley et al., 2005; Houston and Hartley, 2003), a conclusion supported by the end of supergene enrichment of copper deposits in the Atacama Desert (Arancibia and Matthews, 2006). However, this region seems to have had rainfall oscillations throughout its

109

Quaternary history (Betancourt et al., 2000; Latorre et al., 2003). Currently, the mean values of mean annual precipitation recorded between 15°S and 27°S latitude are less than 100 mm/y, for this reason this particular area is considered one of the driest regions on the world (Dillon and Hoffmann, 1997; McKay et al., 2003), and is the object of our principal analysis in this study. Soil carbon concentrations to a specified depth are needed for calculating current soil C stores (Feng et al., 2002; Kern, 1994; Schlesinger et al., 1990). A shift towards a greater area of arid land potentially represents a permanent loss in the productive capacity of the biosphere (Schlesinger et al., 1990). In addition, relating soil carbon to site characteristics may help in formulating and evaluating static and dynamic models of pedogenic processes (Burke et al., 1989), and in assessing the effect of land use and climate change on soil C stores (Bon, 1982; Feng et al., 2002; Grigal and Ohmann, 1992; Plante et al., 2006; Turner et al., 1993). In this context, few analyses have been made showing the relationship between soil carbon and environmental site characteristics for hyper-arid soils, and even for arid–semiarid regions. Indeed, although other authors have calculated the organic or inorganic carbon concentrations in samples of soils from the Atacama Desert (e.g. Navarro-Gonza´lez et al., 2003, 2006; Ewing et al., 2006, 2008; Lester et al., 2007), the importance of the total carbon distribution and its relationship with in situ characteristics have not been completely evaluated, perhaps due to the small number of samples used in those works (Ewing et al., 2008; Ewing et al., 2006). The purpose of this study determines the surface and subsurface soil carbon concentration until one meter deep in the hyper-arid soils of the Atacama Desert in order (a) to evaluate the distribution and deposition of organic and inorganic forms of carbon present there, (b) to compare to arid and semiarid deposits in surrounding areas, (c) to analyze any relationship between carbon concentrations and some geomorphological and climatological variables “in situ”, and (d) to seek differences between the driest areas of Yungay and Pampas de La Joya located in Chile and Peru respectively, which have showed the lowest levels of organic carbon in previous studies (Ewing et al., 2008; Navarro-Gonza´lez et al., 2009, 2006, 2003; Valdivia-Silva, 2009b). Additionally, the predominant abiotic geochemical processes and eolic transportation observed in this region could be used as an excellent analogue in order to understand the carbon cycle present in other hyper-arid environments on Earth and/or Mars. 2. Materials and methods 2.1. Site description The Atacama Desert is located between 10°S and 35°S latitude in Peru and Chile, bounded on the east by the front ranges of the Andes and on the west by the Coast Range. Hyper-arid areas considered in this study were focused

110

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

between 15°S and 27°S latitude because this encompasses the driest areas of the desert. Chilean and Peruvian areas, Yungay region and Pampas de La Joya respectively, present interesting geomorphological differences caused by variations on the tectonic uplift of the desert basin and the adjacent Andean mountains (Figs. 1 and 2). The Chilean intermountain zone, named the Central Depression, shows late Cenozoic fluvial sediment surrounds small hills to rough mountains (1500–3000 m.a.s.l). Between 23°S and 30°S latitude, the desert is dominated by extensive gravelly fluvial deposits, generally referred to as the Atacama Gravels, also probably, originated by fast uplift of the Andes (Gregory-Wodzicki, 2000; Mortimer, 1973) during the Middle to Late Miocene time (Marinovic et al., 1992; Nishiizumi et al., 2005). Subsequent incision into the Atacama Gravels resulted in prominent and extensive, Late Miocene to Pliocene age fans and terraces (Hartley et al., 2005). The Chilean

landforms host well-developed soils rich in nitrate and other exotic salts such as persulfates, iodates, chromates and perchlorates (Ericksen, 1981, 1983). Interestingly, the Peruvian zone does not present high concentrations of nitrates and perchlorates (Valdivia-Silva, 2009b; Valdivia-Silva et al., 2011). The current absence of a mountain range on the coastal side of the Pampas de La Joya (see Fig. 3) and probably throughout the Neogene has been a central factor that limited the topographic isolation and the consequent hyperaridity of this region to shorter intervals, and hence of the longevity and depth of the correspondent basin to generate exotic salts. The Peruvian hyper-arid region in the study area is mainly located in central and southern Peru, between 15°S and 17°S (Fig. 1), and may be defined as a broad geomorphic unit characterized by an uplifted plain limited on the northeast and southwest by the Andean foothills and

Fig. 2. Sampling sites and transversal cross-sections from the desert. The size of the triangles is proportional to the sampling density in the different areas (see Supplementary Table 1S). (A) Peruvian and (B) Chilean regions showing the driest areas of “Pampas de La Joya” and “Yungay” respectively. The differents characteristics between regions are described in the text. H: hyper-arid, A: arid, S: semiarid.

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

111

Fig. 3. Soil organic carbon concentrations and soil bulk densities in the subsurface. Seven pits were dug in the hyper-arid areas of (A) Peru, and (B) Chile. SOC: soil organic carbon, MBD: mean bulk density.

the Cordillera de La Costa respectively (Fig. 2), elevated to an average height of 1200 m.a.s.l., which is divided in several sectors by the spectacular gorges of the Sihuas and Vitor rivers carved across the volcanic and sedimentary cover into the crystalline basement (Fig. 2). In fact, the hilly desert landscape underlain by Precambrian to Paleocene crystalline rocks is characterized by the presence of a discontinuous blanket of white sand (volcanic ashes) derived from the ultraplinian 1600 A.D. eruption of the Huaynaputina volcano (Lavalle´e et al., 2006). The Cordillera de la Costa defines the subdued southern physiographic limit of the Peruvian desert plain (Fig. 3), although not long ago in the Neogene these deposits most probably covered the coastal ranges, indicating relatively recent uplift and exhumation of the crystalline basement, synchronously with the sculpturing of the river gorges. The desert plain in southern Peru may have extended continuously in the past, and its present dissection is the consequence of accelerated uplift processes since the Miocene along the frontal segment of the Andes associated with still active faults, which abruptly cut through the Cenozoic deposits (Hoke et al., 2007; Schildgen et al., 2007). The unconsolidated and semi-consolidated deposits of Quaternary age forming the floor of the desert consist of a few tens of meters of alluvial fan,

fluvial, duricrusts, eolian, and lacustrine beds. True soils are absent, but coarse regolith is ubiquitous, and mostly generated in situ by mechanical weathering due to extreme temperature oscillations that may be up to 40 °C/h between noon and midnight (McKay et al., 2003, 2009; ValdiviaSilva et al., 2011, 2012a). Temperature and precipitation vary with both latitude and elevation within the Atacama Desert. Mean annual temperatures (MAT) range between 16 and 18 °C (mainly varying with elevation and proximity to the coast). Mean annual precipitation (Rochette et al., 2006) based on rain gage data in the region (1000–2000 m above sea level, 15– 30°S) is lesser than 100 mm (Houston and Hartley, 2003; McKay et al., 2003; Valdivia-Silva, 2009b). Marine fog is frequent along the coast at these latitudes, but inland incursion as well as formation of inland radiation fogs, depends on elevation and topographic connection to the coast (Cereceda et al., 2002; Ewing et al., 2006; Rech et al., 2003). In the Atacama Desert, there are vast areas of hyperarid, arid, and semiarid regions covering a total area of 144,000 km2 in Peru, and nearly 220,000 km2 in Chile (Fig. 1). On the basis of our field investigations, the different types of regions were described as hyper-arid, arid, and semiarid, though the hyper-arid region was our

112

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

main objective. The areas of these regions were evaluated using digital maps obtained by GeoMapApp 2.8.2 free software (http://www.geomapapp.org), and the Global Multi-Resolution Topography Synthesis (Ryan et al., 2009), which gave us the percentage of area for each of the 3 types of weather conditions evaluated on the desert. The analysis of the soil area using maps was earlier used by Feng et al. (2002), who evaluated the carbon storage on deserts from China. In that study, photocopies of large-scale maps were cut out, then the pieces of paper were weighed to the nearest 0.1 mg on an analytical balance, and the mass representing known land areas was converted to km2. In our view the direct analysis of areas on digital maps using image analyzer software was much more accurate because the software considered the irregular topography. Importantly, as it was explained above, Atacama Desert encompassed regions with different levels of precipitation and climatic characteristic, therefore it is important define where the area of study is present. Because of this, different studies have shown contradictory results in “hyper-arid” soils of the Atacama but a thorough analysis might show that sampling areas were done in arid and even, semiarid regions. Due to the difficulty of drawing exact boundaries for each type of region, sampling in these areas had a strict relation to precipitation parameters evaluated directly by our weather sensors during periods of 4 y (McKay et al., 2003; Valdivia-Silva et al., 2012a) and the Peruvian and Chilean meteorological stations (METEOCHILE, 2009; SENAMHI, 2008). In addition, we used the data published by Houston and Hartley (2003) and Hartley et al. (2005), whom also determined the different type of climates of the region. All this information together, helped to define the limits of a “hyper-arid” strip along the Peruvian – Chilean coast starting approximately at 500 m up to 2000 m inland, and showing a width between 40 km, at the narrowest part, and about 200 km, at the widest one (Fig. 1) (Fabre´ et al., 2006). Since 0 m up to 500 m inland, the region abruptly changes from semi-humid/semi-arid, in the Pacific shoreline, to arid areas upland, therefore the complexity of the geomorphology in this region can allows the presence of semiarid areas at very short distance from hyper-arid ones as it observed in Fig. 2. 2.2. Sample recollection Between 2005 and 2008, 485, 38, and 21 samples of the surface were collected from hyper-arid, arid, and semiarid areas, respectively from the Atacama Desert (Fig. 2, Table 1 & Table 1S). The triangles showed in Fig. 2 represent the density of sampling in each area where the landscape allows access into the desert. Some characteristics of each sampling area is presented in Table 1S. Additionally, 14, 1, and 1 pits of 1 m depth were dug in the hyper-arid, arid and semiarid region respectively. Data on precipitation, elevation, and current vegetation (when it was present),

were obtained for all sites. For each sample, approximately between 200 and 400 g representing a composite of five individual nearby sites (2 m in radius) were collected from the surface to a depth of 10 cm. The pits were sampled each 10 cm depth using steel cylinders, each with a diameter of 5 cm and a height of 15 cm. The cylinders were pressed vertically into the soil and emptied into sterile polyethylene bags (WhirlpakTM) for transport and storage until analysis. Quadruplicate soil samples were collected at all depths. Due to the high heterogeneity present in these soils, the sampling design used was “stratified random” (Dijkerman, 1981; Einax et al., 1997; Peterson and Calvin, 1996), where our strata are summarized according to the three climatic conditions (hyper-arid, arid and semiarid). Since the focus of this study lies mainly in the hyper-arid area, sampling in this region were also classified according to the main type of landforms found on this desert, such as plains, mountains, evaporite mineral deposits, and gullies or channels (Table 1). 2.3. Organic carbon (SOC) analysis The organic carbon content was evaluated by permanganate titration in acid media as our group has reported before (Fletcher et al., 2012; Navarro-Gonza´lez et al., 2006; Valdivia-Silva et al., 2011). This technique compared to other ones such as calcination and pyrolysis coupled to gas chromatography and mass spectrometry, has proven to be simple, accurate, sensitive, and reproducible for the quantification of labile soil organic carbon in hyper-arid soils (Fletcher et al., 2012; Valdivia-Silva, 2009b; Valdivia-Silva et al., 2011). Since, these soils have shown negligible levels, or almost absence, of recalcitrant carbon (or passive pool), the labile form (or active pool) is the most abundant in these soils and include molecules with biological importance such as aminoacids, nucleotides, lipids, sugars, aliphatic and aromatic hydrocarbons, etc. Importantly, the active pool is the most susceptible to oxidation processes and is the type of organics expected on hyper-arid deserts like Mars (Ewing et al., 2008, 2006; Valdivia-Silva et al., 2009a, 2011). The organic matter content for each 10 cm-layer was the mean and standard deviation from four replicates. Finally, the organic matter content of the entire 1-m topsoil layer was calculated as the sum of the content in each 10 cm considering its respective bulk density. Similar measurements were used for inorganic carbon (SIC). 2.4. Inorganic carbon (SIC) analysis Inorganic carbon in soils includes carbonates and other forms of elemental carbon such as graphite, diamond, coke, and carbon black and requires different techniques to be evaluated. Because elemental analysis obtains the total carbon in the sample, other forms of inorganic carbon can be determined by the simple subtraction of organic and

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

113

Table 1 Sampling sites and characteristics. Region

Hyprearid

Samples (number) Pl Ms Gs Min

Arid Semiarid

234 97 74 80 38 21

Area (km2)

MAP* (mm/ y)

MAT* (°C)

Aridity index

Altitude (m.a.s.l)

Peru

Chile

Total

58,464

100,600

159,064

0.3

18

60.05

1000–2000

NE

NE

200,000

20–30

17–22

0.05–0.2

NE

NE

60

18

0.2–0.5

<1000 and >2000 >2000

H: hyper-arid, A: arid, S: semiarid, Pl: plains, Ms: mountains, Gs: channels and gullies, Min: mineral deposits, NE: no estimated, MAP: mean annual precipitation, MAT: mean annual temperature, m.a.s.l: meters above sea level. * Mean annual values of precipitation and temperature were collected since 2004 up to 2008. Area values from arid and semiarid regions were not assessed due to small number of field data that help to corroborate the limits described by Houston and Hartley (2003).

carbonate carbon. Elemental analysis was done with a model EA1108 analyzer (Fisions, Loughborough, U.K.) at 1200 °C. Carbonate concentration was measured by retro-titration with NaOH and HCl (Navarro-Gonza´lez et al., 2006). Because dolomite and MgCO3 have shown to have minor importance among pedogenic carbonates in these soils (Ewing et al., 2006; Valdivia-Silva, 2009b), in this paper, calcium carbonate equivalent was converted to inorganic carbon by multiplying by 0.12, the mole fraction of carbon in CaCO3. 2.5. Bulk density Bulk density was determined using the known sample volume and the combined oven-dry weight of all sample materials. The detailed analysis method is given elsewhere (Hillel, 1974). 2.6. Statistical analysis The values are expressed as mean ± standard deviation. Independent variables were entered into and retained in the regression models at P < 0.05. In the study, simple linear regression that each parameter is regressed independently against the SOC or SIC was applied (Feng et al., 2002). Comparison analysis were determined by the U Mann Whitney test and analysis of variance [ANOVA] (SAS Institute, 1985). 3. Results and discussion 3.1. Soil organic carbon (SOC) Table 2 and Fig. 3 show the soil organic carbon concentrations in the surface and the first meter of depth from three different climatic regions in the Atacama Desert, respectively. For 485 samples of hyper-arid soils, SOC ranged from 1.8 lg C to 50.9 lg C per g of soil in the 0–0.1 m profile and from 1.8 lg C to 125.2 lg C per g of soil in the 0–1 m layer. Interestingly, the lowest values on surface SOC (2.14 ± 0.8 lg C) were found on the site named “Mar de Cuarzo”, located in the Peruvian region (16° 440

33.3600 S, 72° 20 34.9800 W), which were even lower than those found in Yungay (SOC 10–30 lg C). However, in a general context there were no significant differences between mean organic content values comparing all soil samples from Peruvian and Chilean regions (P = 0.33) (Table 1S & Table 2, and Fig. 3). High variability in surface and subsurface SOC was found in all evaluated landforms (plains, mountains, gullies, and mineral deposits) (Table 2); however, statistical comparisons using the U Mann Whitney test and analysis of variance (ANOVA) did not show significant differences (P > 0.27; P = 0.09, respectively). Since an important point of this study was the determination of SOC contents related with topographical features more clearly exposed to water, the run-off gullies were an important object of our evaluation. Gullies are evident features throughout this region, and typically form starting at about two-thirds of the way up the top of the hills (apparently independent of hill height) and drain into the main valley floors. Rarely the valley drainage systems accumulate much, if any, water, and the run-off in gullies more commonly terminates as soon as they reach the valley floors, with any remaining flow soaking into the sub-surface soils and/or evaporating. These distinct features are unmistakable and, following Gerrard’s premise as well as an apparent greater access to water, the expected results were that there should have been an accumulation of organic materials from summit to toe-slope (Gerrard, 1981). Surprisingly, this was not the case. Gerrard’s premise suggests that there should be an accumulation of biomaterials from summits to toe-slopes due to gravity and water movement, which results in a predictable sequence of soil characteristics. Although, different studies in arid sites from North America and, more importantly, in the hyper-arid Taylor Valley of Antarctica (Burkins et al., 2000; Burkins et al., 2001; Schimel et al., 1985) have demonstrated this premise, landforms examined in this study did not satisfy the predictions, and our data suggested that the Atacama demonstrates no differentiation associated with slope position including those values taken directly within run-off gullies. Some possible explanations about this process are discussed below.

1.4 1.5 1.4 1.6 1.1 1.7 1076.3 ± 430 1150.3 ± 375 1088.7 ± 287 983.7 ± 196 2024.4 ± 414 3082.6 ± 507 Ngl Ngl 2.0 ± 2.5 Ng 320.7 ± 140 512.9 ± 95 1226.7 ± 192 603.8 ± 400 1311.7 ± 248 1150.6 ± 288 2425.0 ± 512 3289.0 ± 334 Ngl Ngl Ngl 3.3 ± 2.1 242.0 ± 91 388.5 ± 104 1027.0 ± 598 733.0 ± 574 1112.7 ± 523 933.4 ± 488 1995.9 ± 214 2757.8 ± 309 20.5 ± 7 28.2 ± 16 27.3 ± 14 33.9 ± 17 120.0 ± 67 465.3 ± 83 26.4 ± 17 32.3 ± 11 31.7 ± 12 35.3 ± 14 87.9 ± 50 420.8 ± 82 Arid Semiarid

Hyerarid

Pl Ms Gs Min

18.1 ± 7 20.2 ± 8 19.0 ± 9 23.1 ± 14 123.5 ± 85 512.2 ± 97

Mean Value Elemental Carbon* Ci

Chile

Elemental Carbon* Chile

Ci

Peru Peru

Mean Value

SIC (lg C/g of soil) SOC (lg C/g of soil) Region

Table 2 Surface soil organic and inorganic carbon concentration (0–10 cm).

* Elemental carbon values were obtained subtracting organic and inorganic carbon values from total carbon assessed by elemental analysis (see Materials and Methods). H: hyper-arid, A: arid, S: semiarid, Pl: plains, Ms: mountains, Gs: channels and gullies, Min: mineral deposits, Ngl: negligible, Ci: inorganic carbon from carbonates, SOC: soil organic carbon, SIC: soil inorganic carbon, MDB: mean build density.

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

MBD (g/m3)

114

As expected, the arid and semiarid regions showed values of surface SOC higher than hyper-arid ones (120.0 ± 67 lg C and 465.3 ± 83 lg C vs. 27.5 ± 5 lg C per g of soil) consistent with precipitation gradients evaluated by other studies (Ewing et al., 2008; Navarro-Gonza´lez et al., 2003). The results of the organic carbon content in the subsurface from hyper-arid soils showed high variability in each horizontal layer. Importantly, these types of profiles are characteristic in dry environments where the genesis of soil horizons depends on the availability of water, atmospheric transportation, and abrupt geologic variations throughout its history (Amundson, 2001; Amundson et al., 2007; Ewing et al., 2006; Valdivia-Silva, 2009b). In general a significant increase related to the depth was observed in hyper-arid soils from 20 lg C on the top up to 152 lg C on the bottom (Fig. 3). The highest values of SOC were evident in the 80–100 cm horizon (>70 lg C), and one pit in the Peruvian region reached the highest value of 202 ± 34.5 lg C per g of soil (Fig. 3A). Curiously, this point belongs to the most northern sampling area (15°S latitude) where the hyper-arid area is the narrowest (Fig. 2A) and humidity gradients are shorter. This behavior in SOC was also evident in the subsurface of arid and semiarid areas (Fig. 5). For the entire data set, storage of soil organic carbon on a whole profile basis averaged 4.2 Tg and 7.2 Tg for Peruvian and Chilean regions respectively in the 0–1 m layer (Table 3). On average, arid–semiarid regions had higher SOC than the hyper-arid one (380.1 Tg vs. 11.5 Tg). The small difference in bulk densities among the different land types had no significant effect on the surface and subsurface organic carbon (r2 = 0.014). According to the results of mean SOC and the areas of different dry land types, the total amount of organic carbon stored in the Atacama Desert was 391.7 Tg in the top 1 m (Table 3) and 105.4 Tg in the first 10 cm depth. This carbon deposit is on average 20 times lower than the organic carbon deposits in the desert of China if we consider the area in both deserts with the same value 360,000 km2 (0.4 Pg vs. 7.84 Pg) (Feng et al., 2002). We used linear regression analysis to relate the dependant variable of SOC profile to independent variables representing characteristics of the different regions: mean annual precipitation (P), mean annual temperature (T), and elevation (E) (Table 4). These relationships were more clearly expressed when the entire data set was subdivided by type of region. Our analysis was restricted to samples for which all of the above variables were known. Among the samples examined, the SOC contents in the 0–1 m profile from hyper-arid, arid and semiarid regions increased significantly with increasing mean annual precipitation (r2 > 0.2, Table 4). The influence of precipitation was more significant on arid and semiarid than hyper-arid regions and this parameter explained more variation in SOC in the 0–1 m profile

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

115

Table 3 Total carbon deposits in soils from the Atacama Desert (1 m layer). Region

TOC (Tg)

Hyperarid Arid and semiarid Total

TIC (Tg)

Total C

Peru

Chile

Subtotal

Peru

Chile

Subtotal

4.29 106.02 110.31

7.27 274.12 281.39

11.56 380.14 391.7

121.92 1456.8 1578.72

222.7 2347.0 2569.7

344.62 3803.8 4148.42

356.18 4183.94 4540.12

TOC: total organic carbon, TIC: total inorganic carbon, Tg: teragram.

Table 4 Linear regressions relating SOC and SIC to average precipitation (P), temperature (T), and altitude (E). Independent variables

SOC 2

H

Plains

Mountains

Gullies

Mineral deposits

A

S

P (mm) T (°C) E (m) P (mm) T (°C) E (m) P (mm) T (°C) E (m) P (mm) T (°C) E (m) P (mm) T (°C) E (m) P (mm) T (°C) E (m)

SIC

(r )

Slope

(r2)

Slope

0.24 0.007 0.001 0.31 0.008 0.009 0.21 0.001 0.001 0.42 0.001 0.004 0.39 0.04 0.004 0.45 0.29 0.28

15.17 0.001 0.000 22.1 0.005 0.002 27.40 0.001 0.002 71.2 0.000 0.001 28.33 0.05 0.000 66.25 0.07 0.17

0.001 0.04 0.001 0.01 0.09 0.001 0.001 0.04 0.000 0.001 0.03 0.000 0.51 0.22 0.001 0.49 0.29 0.14

0.000 0.01 0.000 0.18 0.05 0.000 0.000 0.002 0.000 0.000 0.3 0.000 0.62 0.07 0.000 0.635 0.11 0.38

SOC: soil organic carbon, SIC: soil inorganic carbon.

soil than any other single variable (r2 > 0.38). By contrast, temperature had a negative relationship with SOC on semiarid soils (r2 = 0.29), and no significant or any effect on arid and hyper-arid soils respectively (r2 = 0.04 and r2  0, Table 4). Although temperature has an important relationship with soil organic carbon in different soils (Amundson, 2001), the temperature regime in the Atacama Desert did not show this effect. Generally, SOC increases when the temperature decreases, but only the semiarid region and weakly, the arid region showed this behavior (Table 4). Because elevation is strongly and directly correlated with vegetative biomass and the height of the ground water table (Amundson, 2001), it is not surprising that the SOC contents in the profile also increases significantly with elevation in the semiarid region. However, in the arid and hyper-arid lands, this variable did not show any significant effect (r2  0, Table 4). This result is consistent with the absence of significant differences of organic carbon concentration between the different landforms including mountains or plains located in the hyper-arid region (Table 1). Since the rainfall directly affects soil moisture, the high concentration of labile carbon and even recalcitrant carbon present in semiarid soils can well be explained by the macro

and microscopic biota existent there. Nevertheless, in hyper-arid and arid areas, SOC values are not adequately explained by the relationship with biota, at least on the surface of the desert. Indeed, only rain events greater than 2 mm resulted in detectable liquid water under the stones and on the surface soil (Valdivia-Silva et al., 2012a), and though frequently, fog or dew occurs following high night-time relative humidity, they are not a effective source of moisture in the soil or under stone surfaces (Davis et al., 2010; McKay et al., 2003; Valdivia-Silva, 2009b). In addition, the low level of liquid water was present under stones for only 24–85 h, which is consistent with the low levels of microorganisms in the extreme hyper-arid soils of the Atacama (<104 cells/g of soil) (Drees et al., 2006; Fletcher et al., 2011; Lester et al., 2007; Navarro-Gonza´lez et al., 2003; Warren-Rhodes et al., 2006). Most of the soil bacteria are heterotrophic, and photosynthetic microorganisms are only found within halite crusts, where mineral deliquescence facilitates primary productivity (Davila et al., 2008; Wierzchos et al., 2006). Because this work used “soil” samples, the fluctuations of SOC concentration was not explained by microorganism communities and their relation with precipitation. Indeed, 1 lg of organic C is equivalent to about one million bacteria in terms of carbon

116

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

Fig. 4. Soil inorganic carbon concentrations in the subsurface of the Atacama. Seven pits were dug in the hyperrid areas of (A) Peru, and (B) Chile. SIC: soil inorganic carbon, MBD: mean bulk density.

biomass (Navarro-Gonza´lez et al., 2003; Valdivia-Silva et al., 2012a,b), such that the microorganisms in hyper-arid soils do not seem to have an important contribution to total organic carbon. Importantly, subsurface could be a different scenario. It is demonstrated that soil moisture in different regions of the Atacama can last more time (days) in the subsurface than at the surface (McKay et al., 2003; Valdivia-Silva et al., 2012a). Therefore, the high levels of organic carbon found in the deeper layers of subsoil (Fig. 3) could suggest a more evident relationship between organic carbon due to biomass from microbial communities, and precipitation. More comprehensive microbiological studies in subsurface environments are required in order to demonstrate this affirmation. On the other hand, previous geochemical analyses in these hyper-arid soils have shown high oxidation activity when organic compounds were added to soil samples in aqueous solution (Navarro-Gonza´lez et al., 2003; Peeters et al., 2008; Quinn et al., 2007; Valdivia-Silva et al., 2011). Since water is the key for all geochemical reactions, the low levels of precipitation might be triggering oxidative reactions where oxidant minerals are present, so that it could be a partial explanation to the heterogeneous distribution of organic carbon in these soils (Table 4). Interestingly,

the strongest relationship between SOC and precipitation was found in samples obtained in mineral deposits. Indeed, minimal changes in levels of precipitation affect up to 40% the concentration of C (r2 = 0.42, Table 4). Importantly, although we found correlation between SOC and precipitation in these hyper-arid soils, it was not enough to explain the high variability on the distribution of organic carbon (<40%). Moreover, the absence of Gerard’s premise in the landscape of the Atacama Desert suggested a homogeneous process of distribution of particles. Interestingly, a similar process was found on Mars surface by the Vikings in the 70s, and the “eolic distribution and transport” was considered as the better explanation for it (Margulis et al., 1979). Therefore, the eolic transportation of particles could be leading the distribution of organic matter on the Atacama surface but “in situ” characteristics – like proximity to minerals, clays or salt deposits – and different temporary levels of humidity – like precipitation or fog – could be protecting or destroying the organic matter with the consequent variability on the final distribution and deposit. An interesting study about mineral interaction and organic molecules was published for our group before (Valdivia-Silva et al., 2009a). In that study we showed that soils with more presence of sand preserved more organic molecules in comparison to clays or silts, probably due to particles >50 lm have low chemical reactivity and surface area. It is important to note that the increased presence of minerals in the soil does not necessarily indicate greater amount of oxidants or oxidizing activity (Valdivia-Silva et al., 2012a,b). Certainly many minerals in these soils have not a baseline oxidative property, but require water or a catalyst as temperature in order to manifest such activity. So, although Chilean region has shown a greater quantity of mineral salts in comparison to the Peruvian one, the oxidant activity in these soils could be higher due to a major presence of water by increasing seasonal humidity as it has recently been described for this area (Valdivia-Silva et al., 2012a). Therefore, this process might explain the concentration of organics in soils samples from Pampas de La Joya, and more specifically in the area known as “Mar de Cuarzo” where we found the lowest levels of organics (Valdivia-Silva et al., 2011). On the other hand, arid and semiarid soils showed a better correlation with precipitation, elevation and temperature (Table 4) as it has been described in different studies (Amundson, 2001; Amundson et al., 2006, 2007). 3.2. Soil inorganic carbon (SIC) Based on the mean SIC content of the dry lands studied (Table 2 and Fig. 4), the total amount of inorganic carbon stored on a whole profile basis (1.0 m) averaged 4.1  103 Tg, with 3.4  102 and 3.8  103 Tg stored in hyperarid (over 160,000 km2) and arid–semiarid regions (over 200,000 km2) respectively (Table 3). The top 10 cm of hyper-arid region stores 25.13 Tg of inorganic carbon, whereas the arid–semiarid ones contain 71.4 Tg. The storage

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

117

Fig. 5. Organic and inorganic carbon in the (A) arid, and (B) semiarid regions. SIC content includes elemental carbon and carbon from carbonate minerals. SOC: soil organic carbon, SIC: soil inorganic carbon, MBD: mean bulk density.

of soil inorganic carbon includes carbonate minerals and other forms of elemental carbon; however our results showed negligible amounts of forms of inorganic carbon different to carbonates (Table 2). In addition, although the elemental carbon could contain organic carbon from recalcitrant pools as well (which is not analyzed by our chemical method), it has also been demonstrated that this type of carbon is not present in significant concentrations in hyper-arid soils (Ewing et al., 2006; Valdivia-Silva et al., 2011) as explained in the previous section. Similarly to organic carbon, the SIC concentrations in the surface and subsurface profiles did not show statistical significant differences between the landforms evaluated in the hyper-arid region (P = 0.06) (Table 2), although the Peruvian profiles showed more heterogeneity than the Chilean counterpart (Fig. 4). For 485 samples from hyper-arid soils, the SIC ranged from 570 to 1500 lg C per g of soil in the 0–0.1 m profile and from 570 to 3100 lg C per g of soil in the 0–1 m layer. On the other hand, arid and semiarid region showed major concentrations of SIC ranging from 873 to 3600 lg C per g of soil and from 2700 ± 49 to 10,000 lg C per g of soil in the top 0–10 cm, respectively. Interestingly,

these regions showed the presence of elemental carbon, suggesting other forms of elemental carbon – as graphite – or recalcitrant organic carbon – as roots – undetected by our chemical method. Because elemental analysis cannot differentiate both forms of carbon (recalcitrant organic carbon versus inorganic carbon different from carbonates), the firsts were included as part of the inorganic carbon pool. So, the elemental carbon found on the surface of arid-semiarid areas shows values between 242 ± 91 and 512 ± 95 lg C per g of soil, respectively, whereas that in the top 1 m profile the values increased with the depth up to 2000–3000 lg C per g of soil on the bottom (Fig. 5). Importantly, the variability in the bulk densities among the different types of soil samples had a major effect in the concentration of soil inorganic carbon values (r2 = 0.37, P = 0.04) than that observed over the amount of organic carbon. This relationship between bulk density and carbonates has well been explained by accumulation, water infiltration, evaporation, and precipitation of salt minerals in dry environments (Amundson, 2001; Ewing et al., 2008). On average, carbon stored as inorganic carbon exceeded carbon stored in organic forms by a factor ranging from

118

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

28.4 for Peruvian hyper-arid area, 30.6 for Chilean hyperarid region, and 10 for arid–semiarid soils (Table 3). The mean values of precipitation (<1 mm/y) did not show any relationship with SIC in hyper-arid soils (r2  0, Table 4), while that in the arid and semiarid soils, the SIC content increased strongly with the mean annual precipitation (20–60 mm/y; r2  0.5; Table 4). Importantly, as discussed below, levels of precipitation major than 100 mm/y has an inverse correlation with SIC, as occur in the most deserts wetter than Atacama. Temperature exhibited between 3% and 9% variability of the SIC contents in hyper-arid soils (0.03 < r2 < 0.09) and more than 20% in arid and semiarid soils (r2 = 0.22 and 0.29, Table 4). Because elevation is positively correlated with vegetation biomass and height of ground water tables (Amundson, 2001), it was not surprising that the inorganic carbon content in the semiarid region declined with decreasing elevation (r2 = 0.14), while inorganic carbon concentration in the hyper-arid and arid areas did not present any significant correlation with height or biomass (r2  0). It is known that in many arid soils, the caliche, or calcic horizons, are deposited at a depth related to the maximum infiltration of rainfall (Amundson, 2001; Schlesinger and Pilmanis, 1998). Therefore, the presence of carbonates in hyper-arid soils required the presence of water in the past throughout geologic time, and now these deposits are preserved and/or replaced by sulfates or other salts (Schlesinger and Pilmanis, 1998). Indeed, the low abundance of CaCO3 in Atacama soils has been noted by other authors (Michalski et al., 2004; Rech et al., 2003). While CaCO3 (pKsp = 8.34) is much less soluble than CaSO4 (pKsp = 4.59), its formation is not chemically favored in this hyper-arid setting (Butler, 1982). In general under the environmental conditions of the arid-semiarid lands of the Atacama Desert, while the land is wetter, more inorganic carbon is stored. Nevertheless, this assertion has limits at both ways. In places where there are very low rainfall (<1 mm/y), carbonates are replaced by other salts such as sulfates, and on the contrary, if the rainfall is higher (P100 mm/y), it dissolves, leaches and drags carbonates to subsurface. Additionally, the very low levels of both carbonate and organic C in hyper-arid soils in comparison to aridsemiarid areas are consistent with low biological activity. The relatively high carbonate content in the deep, undisturbed sedimentary strata of this soil is more likely a relict feature (Fig. 4). It coincides with a high increase in organic C (Fig. 3), presence of relict root fragments, and a small apparent weathering loss of Si, Al and Na reported by Ewing et al. (2006). Again, these features suggest that a wetter period preceded to the stabilization of the present landforms approximately 2 My ago during the Quaternary. Since the organic carbon has a short mean residence time in soils and can change rapidly with changes in land-use, inorganic carbon (calcic horizon) typically has a much longer mean residence time and cannot be expected to

change in the short-term. Thus, on a worldwide basis, the carbon pool in carbonate soils and the rate of formation and exchange of carbon between these deposits and the atmosphere require further studies beyond the present, preliminary estimates. 4. Mars-like soils and perspectives The role of the geochemical cycle of carbon into processes which have extremely limited amounts of water on Earth and Mars remain poorly studied and understood. The recent identification of calcium carbonate (3–5 wt.%) in the soils around the Phoenix landing site (Boynton et al., 2009) has increased interest in understanding these processes and has led, as the present study, to seek a better explanation of the carbon geochemistry on hyper-arid soils and environments considered analogues to the Red Planet. On the other hand, the organic carbon storage on Mars regolith also remains unknown and its presence is cause of debate. The Viking mission in the late 1970s conducted a search for organics in the Martian soil but it was unable to detect them on the surface at ppb levels (Biemann, 1979). However, recent studies have shown probable flaws in the method of extraction of organic compounds from the regolith samples by thermal analysis (Navarro-Gonza´lez et al., 2009, 2006, 2010). It is important to emphasize that the work of Navarro-Gonza´lez et al. (2006, 2009a) only raised limitations in the pyrolysis step but no on the GCMS instrument, which they concluded it operated flawlessly as the GC-MS was designed and built (Biemann, 2007; Mukhopadhyay, 2007; Biemann and Bada, 2011). Viking results were especially surprising as even in the absence of biological production or endogenous organic synthesis (Kanavarioti and Mancinelli, 1990), it was estimated that about 2.4  108 g of organic carbon (C) arrives to Mars each year via meteorites (Flynn, 1996). If we consider the accumulation of C over a period of 3 billion years, it is estimated that about 7.2  1017 g of C have been deposited on the Martian surface. The surface of Mars is 144.8  106 km2; if we assume that this C mixed homogenously within the first 1 m layer of the Martian surface, and the density of regolith of 1.37 g/cm3 (Krupenio, 1978), the expected concentration of organics in the Martian surface would be 3650 lg C per g of soil. It was generally believed that such high levels of organics were not detected by the Viking TV-GC-MS (thermo volatilization – gas chromatography – mass spectrometry) because the Martian surface is highly reactive and contains one or more oxidants that converted the organics into carbon dioxide (CO2) over geologic time (McKay et al., 1998; Quinn et al., 2007; Zent and McKay, 1994). Hydrogen peroxide (Atreya et al., 2006; Encrenaz et al., 2005), superoxides (Yen et al., 2000), UV radiation (McKay et al., 1998), peroxide-modified titanium dioxide (Quinn and Zent, 1999), peroxinitrites (Plumb et al., 1989), and recently perchlorates were found by Phoenix spacecraft on Mars surface (Hecht et al., 2009), are possible oxidative candidates.

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

The hyper-arid region of the Atacama Desert presents interesting geomorphological characteristics that together to soils with high oxidant activity, interesting chemical– physical properties, very low levels of organic carbon, and the presence of exotic minerals, including sulfates, chlorates, perchlorates, iodates, and nitrates has been compared to those ones observed on Mars (NavarroGonza´lez et al., 2003; Ewing et al., 2006; Lester et al., 2007; Valdivia-Silva et al., 2011). Moreover, the pedogenic processes in the Chilean hyperarid region provides a powerful terrestrial analog for the accumulation of atmospheric dust and solutes in ancient landscapes on Mars (Amundson et al., 2006; Ewing et al., 2006). Indeed, both vertical redistribution of solutes and dust deposition have been argued with critical implications for Mars climate history (Clark et al., 2005; Yen et al., 2005). Hyper-arid soils from the Atacama Desert have shown the presence of non-chirally specific (Navarro-Gonza´lez et al., 2003; Quinn et al., 2007; Valdivia-Silva et al., 2009a, 2012b) and highly oxidative species, including perchlorates (Ericksen, 1981; Navarro-Gonza´lez et al., 2010) #3490, which do not completely explain the oxidative effect on the organic matter on Mars, but that could explain the variability and extreme very low levels of organic C in these soils, as discussed above. Throughout this study we have shown that all along Atacama Desert, where hyper-arid area is present, there are very low levels of organic and inorganic carbon and not only in Yungay or Pampas de la Joya region where the most studies are being done. Interestingly, the concentration and distribution of carbon were not dependent on the type of geomorphological features (including those could have more water exposure), and they weakly showed relationships with the current environmental conditions analyzed (T, P, E; Table 4). Importantly, specific sites like Mar de Cuarzo that has showed the lowest levels of organic carbon are being more exhaustively studied (Valdivia-Silva et al., 2009a, 2011, 2012b). Since this paper has as objective to estimate the global amount and distribution of carbon in the hyper-arid area, specific relationships between carbon levels and rock mineral composition, type of probable habitable niches, and microenvironments are presented elsewhere. In a general context we showed to this hyper-arid region as a Mars-like soil analogue based on three arguments: (1) the organic carbon is mainly distributed almost homogeneously by eolic transportation like it was observed on particles distribution in Mars regolith (Margulis et al., 1979), (2) the very low levels of organic carbon in all sample taken covering a 160,000 km2 area, and (3) the previous evidence of two distant places 1000 km far each other- Yungay (Navarro-Gonza´lez et al., 2003) and Pampas de La Joya (Valdivia-Silva et al., 2011) – with Mars-like characteristics such as abiotic oxidant activity, geomorphology, extremely low levels of microorganisms, and other geochemistry and geophysical processes. Taking the results of organic C concentration in hyperarid soils of 11.6 Tg (0.012 Pg) and in arid-semiarid soils of

119

380 Tg (0.3 Pg) on the top 1 m profile; the total storage of organic carbon in the Atacama Desert would represent between 0.4 and 0.9% of the pool of organic carbon estimated for the world’s desert soils (42.1–100.8 Pg C), which represents 3–7% of the total organic carbon in the world’s soils (Schlesinger, 1997). Previous studies have shown different analyses for soil organic or inorganic carbon on surface or shallow subsurface from the Atacama Desert (Navarro-Gonza´lez et al., 2003, 2006; Ewing et al., 2006), but most of them were made using a small number of samples, which are not sufficient for predictive models of soil carbon deposits throughout the desert. Because of the need for a more comprehensive analysis, the present study determined the carbon concentrations covering different regions of the desert, mainly focusing on the hyper-arid lands located both in Peru and Chile, and which constitutes an important model for extrapolation of carbon storage in other hyper-arid soils including Mars regolith. Finally, our work provides a framework for interpreting the growing Mars regolith database and for developing testable hypotheses for the presence of organic and inorganic carbon in Mars soils. 5. Conclusions The purpose of this study was to determine the surface and subsurface soil carbon concentration until one meter deep in the hyper-arid soils of the Atacama Desert in order (a) to evaluate the distribution and deposition of organic and inorganic forms of carbon present there, (b) to compare to arid and semiarid deposits in surrounding areas, (c) to analyze any relationship between carbon concentrations and some geomorphological and climatological variables “in situ”, and (d) to seek differences between the driest areas of Yungay and Pampas de La Joya located in Chile and Peru respectively. We conclude that (1) hyper-arid areas present very low levels of SOC between 2 and 50 lg C per g of soil, and values of SIC between 200 and 1500 lg C per g of soil with no significant differences between Peruvian and Chilean region; (2) arid and semiarid areas show high levels of organic carbon directly related to humidity transects; (3) the variability in the concentrations of SOC and SIC in hyper-arid areas has not related to the geomorphological features studied (plains, mountains, gullies, and mineral deposits) although specific areas could have unique characteristics based on mineralogical composition and microenvironments which are not objective in this study; (4) in general the hyper-arid areas do not show statistically significant relationships between the variability of organic and inorganic carbon deposits, and current changes in the temperature, precipitation and elevation. These relationships are more evident in arid and semiarid areas. Finally, unless specific exceptions, the current distribution and carbon accumulation in hyper-arid areas of this desert is mainly addressed by eolic activity and particular properties in certain areas, not measured in this study, which are a good analogue to the

120

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

distribution of particles in Mars and to the possible behavior of carbon on that planet. Acknowledgments Funding for this research comes from Grants from the Universidad Nacional Auto´noma de Me´xico (DGAPA IN107107, IN109110), Consejo Nacional de Ciencia y Tecnologı´a de Me´xico (CONACyT 45810-F, 98466, 121479), fellowship from NASA Postdoctoral Program, and by the National Aeronautics and Space Administration Astrobiology Science and Technology for Exploring Planets Program. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/ j.asr.2012.03.003. References Alpers, C.N., Brimhall, G.H. Middle Miocene climatic change in the Atacama Desert, northern Chile: evidence from supergene mineralization at La Escondida. Geol. Soc. Am. Bull. 100, 1640–1656, 1988. Amundson, R. The carbon budget in soils. Annual Rev Earth Planetary Sci. 29, 535–562, 2001. Amundson, R., Ewing, S., Owen, J., Dietrich, W., Nishiizumi, K., Chadwick, O., Sutter, B., McKay, C. Pedogenic evidence for climate change and aridification on Mars. Geochim. Cosmochim. Acta 70. A15-A15, 2006. Amundson, R., Ewing, S.A., Michalski, G., Thiemens, M., Kendall, C., Nishiizumi, K., McKay, C., Chong, G. The climatic and biotic thresholds on soil elemental cycling along an arid and to hyperarid rainfall gradient. Geochim. Cosmochim. Acta 71. A22-A22, 2007. Arancibia, G., Matthews, S.J., Pe´rez de Arce, C. K–Ar and 40Ar/39Ar geochronology of supergene processes in the Atacama Desert, Northern Chile: tectonic and climatic relations. J. Geol. Soc. 163, 107–118, 2006. Arroyo, M.T.K., Squeo, F.A., Armesto, J.J., Villagran, C. Effects of aridity on plant diversity in the northern Chilean Andes: results of a natural experiment. Ann. Mo. Bot. Garden 75, 55–78, 1988. Atreya, S.K., Wong, A.S., Renno, N.O., Farrell, W.M., Delory, G.T., Sentman, D.D., Cummer, S.A., Marshall, J.R., Rafkin, S.C.R., Catling, D.C. Oxidant enhancement in martian dust devils and storms: implications for life and habitability. Astrobiology 6, 439–450, 2006. Betancourt, J.L., Latorre, C., Rech, J.A., Quade, J., Rylander, K.A. A 22,000-year record of monsoonal precipitation from northern Chile’s Atacama Desert. Science 289, 1542–1546, 2000. Biemann, K. Implications and limitations of the findings of the Viking Organic-analysis experiment. J. Mol. Evol. 14, 65–70, 1979. Biemann, K. On the ability of the Viking gas chromatography–mass spectrometer to detect organic matter. Proc. Natl. Acad. Sci. USA 104, 10310–10313, 2007. Bon, H.L. Estimate of organic carbon in world soils (II). Soil Sci. Soc. Am. J. 46, 1118–1119, 1982. Biemann, K., Bada, J.L. Comment on “Reanalysis of the Viking results suggests perchlorate and organics at midlatitudes on Mars” by Rafael Navarro-Gonza´lez et al. J. Geophys. Res., 116, E12001, doi:10.1029/ 2011JE003869, 2011. Boynton, W.V., Ming, D.W., Kounaves, S.P., Young, S.M.M., Arvidson, R.E., Hecht, M.H., Hoffman, J., Niles, P.B., Hamara, D.K., Quinn,

R.C., Smith, P.H., Sutter, B., Catling, D.C., Morris, R.V. Evidence for calcium carbonate at the Mars phoenix landing site. Science 325, 61– 64, 2009. Burke, I.C., Yonker, C.M., Parton, W.J., Cole, C.V., Flach, K., Schimel, D.S. Texture, climate, and cultivation effects on soil organic matter content in US grassland soils. Soil Sci. Soc. Am. J. 53, 800–803, 1989. Burkins, M.B., Virginia, R.A., Chamberlain, C.P., Wall, D.H. Origin and distribution of soil organic matter in Taylor Valley, Antarctica. Ecology 81, 2377–2391, 2000. Burkins, M.B., Virginia, R.A., Wall, D.H. Organic carbon cycling in Taylor Valley, Antarctica: quantifying soil reservoirs and soil respiration. Global Change Biol. 7, 112–125, 2001. Butler, J.N. Carbon Dioxide Equilibria and Their Applications. AddisonWesley, Reading, MA, 1982. Cereceda, P., Osses, P., Larrain, H., Farı´as, M., Lagos, M., Pinto, R., Schemenauer, R.S. Advective, orographic and radiation fog in the Tarapaca´ region, Chile. Atmos. Res. 64, 261–271, 2002. Clark, B.C., Morris, R.V., McLennan, S.M., Gellert, R., Jolliff, B., Knoll, A.H., Squyres, S.W., Lowenstein, T.K., Ming, D.W., Toscac, N.J., Yeni, A., McSweenp, H.Y., Rieder, R. Chemistry and mineralogy of outcrops at Meridiani Planum. Earth Planet. Sci. Lett. 240, 73–94, 2005. Davila, A.F., Gomez-Silva, B., de los Rios, A., Ascaso, C., Olivares, H., McKay, C.P., Wierzchos, J., Facilitation of endolithic microbial survival in the hyperarid core of the Atacama Desert by mineral deliquescence. Journal of Geophysical Research-Biogeosciences, 113, G01028, 2008. Davis, W.L., de Pater, I., McKay, C.P. Rain infiltration and crust formation in the extreme arid zone of the Atacama Desert, Chile. Planet. Space Sci. 58, 616–622, 2010. Dijkerman, J.C., Field description, morphology and sampling of soils. Course in soil science and water management Agricultural University, Wageningens, Neatherland, 1981. Dillon, M.O., Hoffmann, A.E. Lomas formations of the Atacama Desert, northern Chile, in: Davis, S.D., Heywood, V.H., Herrera-MacBryde, O., Villa-Lobos, J., Hamilton, A.C. (Eds.), Centres of Plant Diversity: The Americas. WWF-IUCN, Cambridge, pp. 528–535, 1997. Drees, K.P., Neilson, J.W., Betancourt, J.L., Quade, J., Henderson, D.A., Pryor, B.M., Maier, R.M. Bacterial community structure in the hyperarid core of the Atacama Desert, Chile. Appl. Environ. Microbiol. 72, 7902–7908, 2006. Dunai, T.J., Gonza´lez-Lo´pez, G.A., Juez-Larre´, J. Oligocene-Miocene age of aridity in the Atacama Desert revealed by exposure dating of erosion-sensitive landforms. Geology 33, 321–324, 2005. Einax, J.W., Zwanziger, H.W., Geiss, S. Chemometrix in Environmental Analysis. VCH - Wiley, EUA, 1997. Encrenaz, T., Bezard, B., Owen, T., Lebonnois, S., Lefevre, F., Greathouse, T., Richter, M., Lacy, J., Atreya, S., Wong, A.S., Forget, F. Infrared imaging spectroscopy of Mars: H2O mapping and determination of CO2 isotopic ratios. Icarus 179, 43–54, 2005. Ericksen, G.E. Geology and Origin of the Chilean Nitrate Deposits, Geological Society Professional Paper 1188. Unites States Government Printing Office, Washington, 1981. Ericksen, G.E. The Chilean nitrate deposits. Am. Sci 71, 366–374, 1983. Evenstar, L., Hartley, A.J., Rice, C., Stuart, F., Mather, A., Chong Dı´az, G., Miocene-Pliocene climate change in the Peru-Chile Desert, in: 6th International Symposium on Andean Geodynamics (ISAG), Barcelona, Spain, pp. 258–260, 2005. Ewing, S.A., Sutter, B., Owen, J., Nishiizumi, K., Sharp, W., Cliff, S.S., Perry, K., Dietrich, W., McKay, C.P., Amundson, R. A threshold in soil formation at Earth’s arid-hyperarid transition. Geochim. Cosmochim. Acta 70, 5293–5322, 2006. Ewing, S.A., Macalady, J.L., Warren-Rhodes, K., McKay, C.P., Amundson, R., Changes in the soil C cycle at the arid-hyperarid transition in the Atacama Desert, Journal of Geophysical Research-Biogeosciences 113, 2008. Fabre´, A., Gauquelin, T., Vilasante, F., Ortega, A., Puig, H. Phosphourus content in five representative landscapes units of the Lomas de Arequipa (Atacama Desert-Peru). Catena 65, 80–86, 2006.

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122 Feng, Q., Endo, K.N., Guodong, C. Soil carbon in desertified land in relation to site characteristics. Geoderma 106, 21–43, 2002. Fletcher, L.E., Conley, C.A., Valdivia-Silva, J.E., Perez-Montan˜o, H.S., Condori-Apaza, R., Kovacs, G.T.A., Glavin, D., McKay, C.P. Determination of low bacterial concentrations in hyper-arid Atacama soils: comparison of biochemical and microscopy methods with realtime Quantitative-PCR. Can. J. Microbiol. 57, 953–963, 2011. Fletcher, L.E., Valdivia-Silva, J.E., Perez-Montan˜o, H.S., CondoriApaza, R., Conley, C.A., McKay, C.P. Variability of organic material in surface horizons of the hyper-arid mars-like soils of the Atacama Desert. Adv. Space Res. 49, 271–279, 2012. Flynn, G.J. The delivery of organic matter from asteroids and comets to the early surface of Mars. Earth Moon Planet. 72, 469–474, 1996. Gerrard, A.J. Soils and Landforms. George, Allen and Unwin, London, 1981. Gregory-Wodzicki, K.M. Uplift history of the central and northern Andes: a review. Geol. Soc. Am. Bull. 112, 1091–1105, 2000. Grigal, D.F., Ohmann, L.F. Carbon storage in upland forests of the lake states. Soil Sci. Soc. Am. J. 56, 938–942, 1992. Hartley, A.J., Chong, G. Late Pliocene age for the Atacama Desert: implications for the desertification of western South America. Geology 30, 43–46, 2002. Hartley, A.J., Chong, G., Houston, J., Mather, A.E. 150 million years of climatic stability: evidence from the Atacama Desert, northern Chile. J. Geol. Soc. 162, 421–424, 2005. Hecht, M.H., Kounaves, S.P., Quinn, R.C., West, S.J., Young, S.M.M., Ming, D.W., Catling, D.C., Clark, B.C., Boynton, W.V., Hoffman, J., DeFlores, L.P., Gospodinova, K., Kapit, J., Smith, P.H. Detection of perchlorate and the soluble chemistry of Martian soil at the phoenix lander site. Science 325, 64–67, 2009. Hillel, D., Methods of laboratory and field investigation of physical properties of soils, in: Trans. 10th Int. Soil Science Congress, Moscow, pp. 301–308, 1974. Hoke, G.D., Isacks, B.L., Jordan, T.E., Blanco, N., Tomlinson, A.J., Ramezani, J., Geomorphic evidence for post-10 Ma uplift of the western flank of the central Andes 18°300 -22°S, Tectonics 26, TC5021, 2007. Houerou, H.N.L. Climate change, drought, and desertification. J. Arid Environ. 34, 133–185, 1996. Houston, J., Hartley, A.J. The Central Andean West-slope rainshadow and its potential contribution to the origen of hyperaridity in the Atacama Desert. Int. J. Climatol. 23, 1453–1464, 2003. Kanavarioti, A., Mancinelli, R.L. Could organic matter have been preserved on Mars for 3.5 billion years? Icarus 84, 196–200, 1990. Kern, J.S. Spatial patterns of soil organic carbon in the contiguous United States. Soil Sci. Soc. Am. J. 58, 438–448, 1994. Kottek, M., Grieser, J., Beck, C., Rudolf, B., Rubel, F. World map of Koppen-Geiger climate classification updated. Meteor. Z 15, 259–263, 2006. Krupenio, N.N. An estimate of the density of Martian soil from radiophysical measurements in the 3-cm region. Cosm. Res. 16, 357– 365, 1978. Latorre, C., Betancourt, J.L., Rylander, K., Quade, J., Matthei, O. A vegetational history from the arid prepuna of northern Chile (22R– 23RS) over the last 13,500 years. Palaeogeogr. Palaeoclimatol. Palaeoecol. 194, 223–236, 2003. Lavalle´e, Y.L., de Silva, S., Salas, G., Byrnes, J.M., Explosive volcanism without caldera formation: insight from Huaynaputina volcano, southern Peru: Bulletin Volcanologique 68, 333–348, 2006. Lester, E.D., Satomi, M., Ponce, A. Microflora of extreme arid Atacama Desert soils. Soil Biol. Biochem. 39, 704–708, 2007. Margulis, L., Mazur, P., Barghoorn, E.S., Halvorson, H.O., Jukes, T.H., Kaplan, I.R. Viking mission – implications for life on Mars. J. Mol. Evol. 14, 223–232, 1979. Marinovic, N., Smoje, I., Maksaev, V., Herve´, M., Mpodozis, C., Hoja Aguas Blancas, Regio´n de Antofagasta, Carta Geolo´gica de Chile. Servicio Nacional de Geologı´a y Minerı´a, Santiago, Chile, 1992.

121

McKay, C.P., Grunthaner, F.J., Lane, A.L., Herring, M., Bartman, R.K., Ksendzov, A., Manning, C.M., Lamb, J.L., Williams, R.M., Ricco, A.J., Butler, M.A., Murray, B.C., Quinn, R.C., Zent, A.P., Klein, H.P., Levin, G.V. The Mars Oxidant experiment (MOx) for Mars ‘96. Planet. Space Sci. 46, 769–777, 1998. McKay, C.P., Friedmann, E.I., Gomez-Silva, B., Caceres-Villanueva, L., Andersen, D.T., Landheim, R. Temperature and moisture conditions for life in the extreme arid region of the Atacama Desert: Four years of observations including the El Nino of 1997–1998. Astrobiology 3, 393– 406, 2003. McKay, C.P., Molaro, J.L., Marinova, M.M. High-frequency rock temperature data from hyper-arid desert environments in the Atacama and the Antarctic Dry Valleys and implications for rock weathering. Geomorphology 110, 182–187, 2009. METEOCHILE, Boletin de informacio´n meteorolo´gica, Direccio´n Meteorolo´gica de Chile www.meteochile.cl, Santiago, Chile, 2009. Michalski, G., Bohlke, J.K., Thiemens, M. Long term atmospheric deposition as the source of nitrate and other salts in the Atacama desert, Chile: new evidence from mass-independent oxygen isotopic compositions. Geochemica et Cosmochemica Acta 68, 4023–4038, 2004. Mortimer, C. The Cenozoic history of the southern Atacama Desert, Chile. J. Geol. Soc. 129, 505–526, 1973. Mukhopadhyay, R. The Viking GC/MS and search for organics on Mars. Anal. Chem. 79, 7249–7256, 2007. Navarro-Gonza´lez, R., Rainey, F.A., Molina, P., Bagaley, D.R., Hollen, B.J., de la Rosa, J., Small, A.M., Quinn, R.C., Grunthaner, F.J., Caceres, L., Gomez-Silva, B., McKay, C.P. Mars-like soils in the Atacama Desert, Chile, and the dry limit of microbial life. Science 302, 1018–1021, 2003. Navarro-Gonza´lez, R., Navarro, K.F., de la Rosa, J., Iniguez, E., Molina, P., Miranda, L.D., Morales, P., Cienfuegos, E., Coll, P., Raulin, F., Amils, R., McKay, C.P. The limitations on organic detection in Marslike soils by thermal volatilization-gas chromatography-MS and their implications for the Viking results. Proc. Natl. Acad. Sci. USA 103, 16089–16094, 2006. Navarro-Gonza´lez, R., In˜iguez, E., de la Rosa, J., McKay, C.P., Characterization of Organics, Microorganisms, Desert Soils, and Mars-like Soils by Thermal Volatilization Coupled to Mass Spectrometry and Their Implications for the Search for Organics on Mars by Phoenix and Future Space Missions. Astrobiology, 703–715, 2009. Navarro-Gonza´lez, R., Vargas, E., de la Rosa, J., Raga, A., McKay, C.P. Reanalysis of the Viking results suggests perchlorate and organics at mid-latitudes on Mars. J. Geophys. Res. – Planets 115, E12010, 2010. Nishiizumi, K., Caffee, M.W., Finkel, R.C., Brimhall, G., Mote, G. Remnants of a fossil alluvial fan landscape of Miocene age in the Atacama Desert of northern chile using cosmogenic nuclide exposure age dating. Earth Planet. Sci. Lett. 237, 499–507, 2005. Peeters, Z., Quinn, R., Martins, Z., Becker, L., Brucato, J., Willis, P., Grunthaner, F., Ehrenfreund, P., Mars regolith analogues – interactions between mineralogic and organic compounds 39th Lunar and Planetary Science Conference, (Lunar and Planetary Science XXXIX). LPI Contribution No. 1391. p. 1742, League City, Texas, p. 1742, 2008. Peterson, R.J., Calvin, L.D., Sampling, in: Sparks, N. (Ed.), Methods of soil analysis. Part 3. SSSA, Wisconsin, EUA, 1996. Plante, A.F., Conant, R.T., Stewart, C.E., Paustian, K., Six, J. Impact of soil texture on the distribution of soil organic matter in physical and chemical fractions. Soil Sci. Soc. Am. J. 70, 287–296, 2006. Plumb, R.C., Tantayanon, R., Libby, M., Xu, W.W. Chemical-model for Viking Biology Experiments – implications for the composition of the Martian regolith. Nature 338, 633–635, 1989. Quinn, R.C., Zent, A.P. Peroxide-modified titanium dioxide: a chemical analog of putative Martian soil oxidants. Orig. Life Evol. Biosph. 29, 59–72, 1999. Quinn, R.C., Ehrenfreund, P., Grunthaner, F.J., Taylor, C.L., Zent, A.P., Decomposition of aqueous organic compounds in the Atacama Desert and in Martian soils. Journal of Geophysical Research-Biogeosciences, 112, G04S18, 2007.

122

J.E. Valdivia-Silva et al. / Advances in Space Research 50 (2012) 108–122

Rech, J.A., Quade, J., Hart, W.S. Isotopic evidence for the source of Ca and S in soil gypsum, anhydrite and calcite in the Atacama Desert, Chile. Geochim. Cosmochim. Acta 67, 575–586, 2003. Rochette, P., Gattacceca, J., Chevrier, V., Mathe, P.E., Menvielle, M., Mappa Sci, T. Magnetism, iron minerals, and life on Mars. Astrobiology 6, 423–436, 2006. Ryan, W.B.F., Carbotte, S.M., Coplan, J.O., O’Hara, S., Melkonian, A., Arko, R., Weissel, R.A., Ferrini, V., Goodwillie, A., Nitsche, F., Bonczkowski, J., Zemsky, R. Global multi-resolution topography synthesis. Geochem. Geophys. Geosyst. 10, Q03014, 2009. SAS Institute, SAS/STAT guide for personal computers. Version 6. SAS Institute, Cary, NC, 1985. Schildgen, T.F., Hodges, K.V., Whipple, K.X., Reiners, P.W., Pringle, M.S. Uplift of the western margin of the Andean plateau revealed from canyon incision history, southern Peru. Geology 35, 523–525, 2007. Schimel, D.S., Coleman, D.C., Horton, K.A. Soil organic matter dynamics in paired rangeland and cropland topo-sequences in North Dakota. Geoderma 36, 201–214, 1985. Schlesinger, W.H., Reynold, J.F., Cunningham, G.L., Huenneke, L.F., Jarrell, W.M., Virginia, R.A., Whitford, W.G. Biological feedbacks in global desertification. Science 247, 1043–1048, 1990. Schlesinger, W.H. Science and the desertification debate. J. Arid Environ. 37, 599–608, 1997. Schlesinger, W.H., Pilmanis, A.M. Plant-soil interactions deserts. Biogeochemistry 42, 169–187, 1998. SENAMHI, Informacio´n del Tiempo, Clima y Agua, Servicio Nacional de Metereologı´a e Hidrologı´a del Peru´, http//www.senamhi.gob.pe, Lima, Peru´, 2008. Thornwaite, C.W. An approach toward a rational classification of climate. Geogr. Rev. 38, 55–94, 1948. Turner, D.P., Lee, J.J., Koerper, G.J., Barker, J.R., The forest sector carbon budget of the United States. Carbon pools and flux under alternative policy options. USEPA Rep. 600/3-93/093, USEPA, Corvallis, OR, 1993. UNEP, World atlas of desertification (2°edition), in: Programme, U.N.E. (Ed.), Nairobi, p. 182, 1997. Valdivia-Silva, J.E., Navarro-Gonzalez, R., McKay, C. Thermally evolved gas analysis (TEGA) of hyperarid soils doped with microorganisms from the Atacama Desert in southern Peru: implications for the Phoenix mission. Adv. Space Res. 44, 254–266, 2009a.

Valdivia-Silva, J.E., Chemical and Microbiological study of the La Joya Desert, Arequipa-Peru, as a new possible analogue to Mars (PhD Thesis), Postgraduate in Biology Sciences. Universidad Nacional Autonoma de Mexico, p. 261, Mexico, p. 261, 2009. Valdivia-Silva, J.E., Navarro-Gonza´lez, R., Ortega-Gutierrez, F., Fletcher, L.E., Perez-Montan˜o, H.S., Condori-Apaza, R., McKay, C.P. Multidisciplinary approach of the hyperarid desert of Pampas de La Joya in southern Peru as a new Mars-like soil analogue. Geochim. Cosmochim. Acta 75, 1975–1991, 2011. Valdivia-Silva, J.E., Navarro-Gonza´lez, R., Fletcher, L.E., Pe´rez-Montan˜o, S., Condori, R., Ortega-Gutierrez, F., McKay, C.P. Climatological characteristics in the extreme hyper-arid region of Pampas de La Joya, Peru. Astrobiological approach in four years of observation: 2004–2008. Int. J. Astrobiol. 11, 25–35, 2012a. Valdivia-Silva, J.E., Navarro-Gonza´lez, R., de la Rosa, J., McKay, C.P. Decomposition of Sodium formate and L- and D- alanine in the Pampas de La Joya soils: Implications as a new geochemical analogue to Martian regolith. Adv. Space Res. 49, 821–833, 2012b. Warren-Rhodes, K.A., Rhodes, K.L., Pointing, S.B., Ewing, S.A., Lacap, D.C., Gomez-Silva, B., Amundson, R., Friedmann, E.I., McKay, C.P. Hypolithic cyanobacteria, dry limit of photosynthesis, and microbial ecology in the hyperarid Atacama Desert. Microb. Ecol. 52, 389–398, 2006. Wierzchos, J., Ascaso, C., McKay, C.P. Endolithic cyanobacteria in halite rocks from the hyperarid core of the Atacama Desert. Astrobiology 6, 415–422, 2006. Yen, A., Gellert, R., Schroder, C., Morris, R.V., Bell III, J.F., Knudson, A.T., Clark, B.C., Ming, D.W., Crisp, J.A., Arvidson, R.E., Blaney, D., Bru¨ckner, J., Christensen, P.R., DesMarais, D.J., de Souza Jr, P.A., Economou, T.E., Ghosh, A., Hahn, B.C., Herkenhoff, K.E., Haskin, L.A., Hurowitz, J.A., Joliff, B.L., Johnson, J.R., Klingelho¨fer, G., Madsen, M.B., McLennan, S.M., McSween, H.Y., Richter, L., Rieder, R., Rodionov, D., Soderblom, L., Squyres, S.W., Tosca, N.J., Wang, A., Wyatt, M., Zipfe, J. An integrated view of the chemistry and mineralogy of Martian soils. Nature 436, 49–54, 2005. Yen, A.S., Kim, S.S., Hecht, M.H., Frant, M.S., Murray, B. Evidence that the reactivity of the Martian soil is due to superoxide ions. Science 289, 1909–1912, 2000. Zent, A.P., McKay, C.P. The chemical-reactivity of the Martian soil and implications for future missions. Icarus 108, 146–157, 1994.