Modern Meteoric 36Cl Deposition in the Atacama Desert, Chile

Modern atmospheric deposition along a west‐east transect in the Atacama Desert, Chile, was collected to constrain the meteoric 36Cl/Cl ratio and 36Cl deposition flux. The 36Cl deposition flux had a threefold change, from 3.62(±0.18) to 11.6(±0.2) atoms m−2 s−1, going from the coast to the high Andes. The 36Cl deposition was mainly as dry deposition, and the magnitude and differences between sites were attributed to stratosphere‐troposphere transport and chloride deficits by acid displacement, respectively. The meteoric 36Cl/Cl ratios varied greatly from 31.5(±1.1) × 10−15 to 247(±10) × 10−15, which was attributed to local inputs of oceanic chloride near the coast or chloride minerals entrained from nearby salt playas. This study presented refined estimates of the 36Cl deposition flux and meteoric 36Cl/Cl ratio in a desert region in southern tropical zone. The data set will provide a baseline for using natural 36Cl abundances to date salt accumulation in the Atacama.


Introduction
Meteoric 36 Cl/Cl ratios and 36 Cl deposition fluxes to the surface are important biogeochemical tracers because they can be used as chloride source identifiers, to trace hydrologic flow, or as geochronometers. Approximately two thirds of meteoric 36 Cl is produced in the stratosphere through the spallation of 40 Ar by primary and secondary cosmic rays (Davis & Schaeffer, 1955;Lal & Peters, 1967;Phillips, 2000), which is then transported to the troposphere during episodes of stratosphere-troposphere transport (STT) (Bentley & Davis, 1982;Lal & Peters, 1967;Masarik & Beer, 1999). In the troposphere, meteoric 36 Cl mixes with stable chlorine, and they are scavenged to the surface primarily by wet deposition or more slowly via dry deposition. Therefore, meteoric 36 Cl/Cl ratios and 36 Cl deposition fluxes can be determined by 36 Cl production, STT, and local scavenging schemes (Lal & Peters, 1967;Phillips, 2000). Due to its long half-life (301 Kyr) and conservative behavior, meteoric 36 Cl has been successfully used to assess atmospheric mixing (Blinov et al., 2000;Johnston & McDermott, 2008), discern chloride sources (Rao et al., 2005;Sherif et al., 2019), and determine ages of groundwater, soils, evaporites, and ice cores (Cartwright et al., 2017;Jannik et al., 1991;Liu et al., 1994;Phillips, 2000;Thompson et al., 1997). Knowledge of worldwide meteoric meteorites/ice or detect " 36 Cl bomb pulses" generated by thermonuclear tests in the 1950s (Elmore et al., 1982;Finkel et al., 1980;Nishiizumi et al., 1993;Phillips et al., 1990;Suter et al., 1987;Synal et al., 1990). Later, 36 Cl was measured in precipitation of Israel (Herut et al., 1992), the United States (Hainsworth et al., 1994;Knies et al., 1994), Australia (Keywood et al., 1998), Japan (Tosaki et al., 2012), across the Europe (Johnston & McDermott, 2008;Pupier et al., 2016;Santos et al., 2004), and at other places around the planet (Scheffel et al., 1999). Moysey et al. (2003) explored the nationwide 36 Cl/Cl ratios of meteoric groundwater to show the spatial patterns of meteoric 36 Cl/Cl ratio across the United States. However, on the whole, meteoric 36 Cl measurements have been limited, and discrepancies remain between measurements and model estimates. The latter have been obtained by coupling the theoretical calculation of atmospheric production and deposition flux as a function of precipitation (Blinov et al., 2000;Keywood et al., 1998;Phillips, 2000;Pupier et al., 2016;Scheffel et al., 1999). Moreover, the constraints on meteoric 36 Cl/Cl ratios or 36 Cl deposition fluxes have been mostly studied in middle-to high-latitude regions of northern hemisphere, with only two studies in tropical regions or southern hemisphere by Keywood et al. (1998) and Scheffel et al. (1999). Further, the 36 Cl measurements were mainly conducted on precipitation samples, and only a few studies measured the dry deposition of 36 Cl, such as in the continental United States where it accounted for~1/4 of the total 36 Cl deposition (Hainsworth et al., 1994;Moysey et al., 2003). However, in desert regions where the role of wet deposition diminishes and dry deposition becomes more important, there have been many studies that have applied meteoric 36 Cl (see Phillips, 2000), yet the 36 Cl deposition constraints are rare.
This study examines the spatial distributions of 36 Cl deposition flux and meteoric 36 Cl/Cl ratio across a west-east transect in the Atacama Desert, Chile. The Atacama has widespread salt deposits that are mainly attributed to long-term accumulation of atmospheric deposition under the hyperarid environment (Ewing et al., 2007;Sturchio et al., 2009). Significant chloride accumulation has likely occurred throughout the period of hyperaridity in the Atacama, but there is little consensus on when hyperaridity began. One way to date chloride accumulation is to use meteoric 36 Cl with known meteoric 36 Cl/Cl ratios or 36 Cl deposition fluxes, which would have implications for the aridity history of the Atacama. The Atacama's hyperaridity suggests that salt accumulation by dry deposition is significantly more important than wet deposition, making the Atacama unique relative to most regions where wet deposition dominates. In addition, the Andes has been reported to be a region with uniquely intense STT (Škerlak et al., 2014), some of the highest in the world. Further, despite its remoteness, the Atacama is not free from anthropogenic activities (Wang et al., 2014), but how anthropogenic activities have impacted salt deposition, including chloride, is inadequately investigated. This study furthers our understandings of 36 Cl deposition in southern hemisphere tropical zone and in hyperarid environments, where data are limited and there are interesting meteoric 36 Cl applications.

Methods and Material
The Atacama Desert, Chile, stretches more than 1,000 km along the Pacific coast of South America (20-30°S). The Atacama Desert consists of three major physiographic units from west to east: the Coastal Range, the Central Depression, and the Andes (Ericksen, 1981) (Figure 1). The Coastal Range stretches~3,000 km along the Pacific coast; the Andes comprises a series of plateaus (~4,000 m in altitude) and a succession of parallel pre-Andean mountain ranges with intervening valleys and basins; the Central Depression is an intermediate depression with minimal precipitation (<2 mm yr −1 ) extending from~22°S to 26°S (Houston, 2006). Salt playa ("salar" in Spanish), formed by drainages from large basins and subsequent evaporation, spread over~7,250 km 2 in northern Chile as prominent features (Stoertz & Ericksen, 1974).
Atmospheric deposition was collected using an array of nine dust traps set up along a west-east transect across the Atacama ( Figure 1) and left exposed from 10 July 2007 to 1 January 2010 (Wang et al., 2014). Each trap (T) was identified as T1 on the Coastal Range, T2-T6 in the Central Depression, and T7-T10 on the Andes (T7: the pre-Andean range, T8: the Atacama Basin, and T10: the Andean Plateau). The sampling transect encompassed the rims of Salar de Pampa Blanca and Salar de Atacama (Figure 1). Each trap consisted of a single-piece Bundt cake pan (outer ring diameter: 25 cm, surface area: 0.05 m 2 ) fitted with a circular piece of 1/4-inch-mesh galvanized screen on which a layer of pre-washed glass marbles was suspended to mimic desert pavement surfaces (Reheis et al., 1995). The traps were mounted on~1-m-high poles above the ground to eliminate most saltating particles. Atmospheric deposition samples were recovered by washing the pans, marbles, and mesh with Millipore™ water (resistivity: >18 MΩ·cm) into 1-L plastic bottles. The washing solutions were kept frozen before being shipped to Purdue University overnight and gradually freeze-dried in the lab. In addition, one sample from the Salar de Pampa Blanca was collected to constrain the 36 Cl/Cl ratios of surface minerals in salt playas. The resulting bulk dust and salar samples were washed with 30-ml Millipore™ water, filtered through a 0.45-μm polytetrafluoroethylene filter to separate the insoluble and soluble fractions. One 3-ml aliquot of the filtrate was analyzed for Na + and Cl − concentrations with <5% uncertainties by Thermo Scientific iCAP 6500 inductively coupled plasma-optical emission spectroscopy (ICP-OES) and Dionex DX-500 ion chromatography (IC), respectively. The ICP-OES and IC data were used to calculate ion deposition rates that were previously reported in Wang et al. (2014).
Based on the measured Cl − concentrations, another split of each sample was withdrawn from the filtrate and a suitable amount of 36 Cl-free chloride carrier was added to obtain enough Cl − for the 36 Cl/Cl (Cl: total chlorine, mainly 35 Cl + 37 Cl) measurement (Table 1). Ba(NO 3 ) 2 salt was added to the solution to eliminate SO 4 2− as BaSO 4(s) since 36 S could interfere 36 Cl measurements. Excess AgNO 3 salt was then added to precipitate AgCl (s) , which was then isolated by centrifuging and filtration, washed five times in Millipore™ water, and freeze-dried. The purified AgCl (s) was loaded into a copper target and analyzed by AMS at the Purdue Rare Isotope Measurement (PRIME) laboratory according to the general methodology by Sharma et al. (2000). Dilution of a NIST 36 Cl standard (SRM 4422L) with the calibrated 36 Cl/Cl ratio of 1.537 × 10 −10 was used for normalization. Six AMS runs were performed on each sample with precision generally ≤5%. Procedural blanks prepared along with the samples had 36 Cl/Cl ratios ≤6.9 × 10 −15 , and all reported values were blank-corrected (Table 1). The stated error for AMS-measured 36 Cl/Cl ratio represented 1 standard deviation that was computed as a weighted average of internal (e.g., counting statistics and beam fluctuations) and external (e.g., replicates, standards, and blanks) errors. Meteoric 36 Cl/Cl ratios were obtained after the correction of AMS-measured 36 Cl/Cl by subtracting the carrier chloride from the total chloride (Table 1). The 36 Cl deposition flux was calculated by multiplying the meteoric 36 Cl/Cl ratios and chloride deposition rates previously reported in Wang et al. (2014) (Table 1).

Discussion
We hypothesize that the eightfold change in meteoric 36 Cl/Cl ratio and the threefold change in 36 Cl deposition flux over a relatively narrow geographic transect cannot be attributed to variations in meteoric 36 Cl  production, but variations in chloride sources and local atmospheric processing, respectively. This is because the solar activity and interplanetary (or Earth's) magnetic field that together determine 36 Cl production in the atmosphere can be considered invariant along our~320-km-wide sampling transect. The ocean and salt playas are two major terrestrial chloride sources (Li et al., 2019;Moysey et al., 2003;Wang et al., 2014), and the terrestrial chlorides typically have very low 36 Cl activities because of the 1-2 orders of magnitude slower in situ 36 Cl production compared to meteoric 36 Cl production and the non-stop radioactive decay of 36 Cl (Fifield et al., 2013;Phillips, 2000). Therefore, inputs of terrestrial chloride of very low 36 Cl activities can potentially lower meteoric 36 Cl/Cl ratios (Moysey et al., 2003). The 36 Cl deposition flux, on the other hand, should be relatively insensitive to local terrestrial sources but mainly impacted by atmospheric processes such as STT, and chloride displacement by atmospheric acids (Wang et al., 2014).

Distribution of 36 Cl Deposition Flux
The largest 36 Cl deposition flux (11.6 ± 0.2 atoms m −2 s −1 ) at the T10 site is about two times more than the 36 Cl deposition flux estimated by the model by Phillips (2000). The estimated 36 Cl deposition flux (F 36 (P, λ)) mainly depends on 36 Cl production, and precipitation rate: where F 36 (λ) is the average 36 Cl deposition flux as a function of latitude (λ), S F (λ 0 ) is the slope of the 36 Cl deposition flux versus precipitation rate determined at the latitude λ 0 , F 36 (λ 0 ) is the average 36 Cl deposition flux at the latitude λ 0 , P is the precipitation rate at the study site, and P λ ð Þ is the average precipitation rate for the latitude (λ) belt. Since T10 is at 23.10°S, F 36 (λ) was taken as 22.5 atoms m −2 s −1 from Phillips (2000) and P λ ð Þ was taken as 795 mm yr −1 from Baumgartner and Reichel (1975). P near the T10 site during 2007-2009 was reported to be 50 mm yr −1 as snow (DGA, 2010). S F (λ 0 )/F 36 (λ 0 ) was taken as 1.11 × 10 −3 yr mm −1 based on Phillips (2000) using data mainly from Knies et al. (1994), Hainsworth et al. (1994), and Keywood et al. (1998). This yielded an estimated 36 Cl deposition flux of 3.89 atoms m −2 s −1 at the T10 site, three times less than the measurement (11.6 ± 0.2 atoms m −2 s −1 ).
We suggest that the discrepancy between the modeled and measured 36 Cl deposition flux at the T10 site may be because of the neglection of strong STT around the Andes. The Andean range induces significant STT (Škerlak et al., 2014) due to upper tropospheric turbulence caused by breaking mountain waves (Lamarque et al., 1996;Worthington, 1998). Using the ERA-Interim reanalysis data set, Škerlak et al. (2014) compiled a global climatology of STT from 1979 to 2011 to find seasonal STT hotspots over the Andes. During September to February, the STT air mass flux was on the order of 500 kg km −2 s −1 over the Andes, the highest in southern hemisphere and among the highest on Earth. Further, due to tall boundary layer heights over tropical deserts (von Engeln & Teixeira, 2013), the Andes region had the most "deep STT events" in southern hemisphere, during which stratospheric air reaches the boundary layer (Škerlak et al., 2014). Stratospheric 36 Cl should be predominantly in the form of HCl (g) (Wahlen et al., 1991) due to the highly acidic nature of stratospheric aerosols (cf. Kremser et al., 2016). Stratospheric 36 Cl carried into the troposphere during STT events could react on existing aerosols such as mineral dust (Tobo et al., 2009), coat aerosol surfaces (externally mixed), absorb into deliquesced aerosols (Wang et al., 2015), or exchange with existing chloride aerosols (Lal & Peters, 1967), before depositing to the surface. Therefore, a higher 36 Cl deposition flux at the T10 site than the model estimate based on Equation 1 may be due to the significant regional STT that is not accounted for by the models of Phillips and others (Lal & Peters, 1967;Parrat et al., 1996;Phillips, 2000).
The 36 Cl deposition fluxes at the T1-T8 sites, ranging from 3.62(±0.18) to 7.01(±0.16) atoms m −2 s −1 , were 39-69% lower than at the T10 site (Table 1), which could be partially due to the lack of precipitation. During 2007-2009, fog deposition of~22 L m −2 yr −1 (~22 mm yr −1 of precipitation equivalent) at the T1 site (Wang et al., 2014) and rainfall of~10 mm yr −1 at T7 and T8 sites occurred, while no precipitation was recorded at the T2-T6 sites (DGA, 2010). Using Equation 1, the 36 Cl deposition flux was estimated to be 2.64-3.19 atoms m −2 s −1 at the T1-T8 sites, about half of the measured 3.62(±0.18)-7.01(±0.16) atoms m −2 s −1 (Table 1), supporting the hypothesis of elevated regional STT. However, the model estimates for the T1-T8 sites are only slightly less than that for the T10 site, suggesting that precipitation alone cannot account for the significantly lower 36 Cl deposition flux measurements at the T1-T8 sites relative to T10 site. Instead, a correlation between 36 Cl deposition fluxes and chloride deficits (Figure 2a) might offer a mechanistic explanation for the differences between model estimates and measurements.
When aerosol chloride reacts with atmospheric acids such as HNO 3(g) and H 2 SO 4(g) , the more volatile acid HCl (g) is liberated resulting in the loss of chloride (both 36 Cl and stable chlorine) from aerosols (Newberg et al., 2005). This was often referred to as a "chloride deficit" and commonly observed in aerosols along the coastal regions near large cities (Harkel, 1997;Newberg et al., 2005). Our previous work indicated that sea-salt aerosols (SSAs) and surface halite (NaCl) minerals were the dominant sources of chloride at the T1 and T2-T8 sites, respectively (Wang et al., 2014). The initial Cl − /Na + molar ratio, prior to acid replacement reactions, is 1.17 at the T1 site as SSAs (the same as seawater) or 1.0 at the T2-T8 sites as halite mineral. The chloride deficits (%Cl deficit ) were then calculated as {1 − Cl − /[Na + × (Cl − /Na + ) initial ]} × 100%, achieving 28.8% to 85.2% for the T1-T10 sites (Table 1) as reported in Wang et al. (2014). Because of Na + /(Cl − + NO 3 − ) of 1.0 ± 0.1 at the T2-T8 sites (Wang et al., 2014), HNO 3(g) + NaCl (s) ➔ NaNO 3(s) + HCl (g) could be the dominant chloride deficit mechanism. This HNO 3(g) was photochemically produced from anthropogenic NO x generated by transport and power/chemical plants, which was confirmed by a high 17 O excess with Δ 17 O of~28‰ (Ewing et al., 2007;Michalski et al., 2003;Wang et al., 2014).
The effect of the chloride deficit was then accounted for in the model estimation of 36 Cl deposition flux. First, the average difference in the modeled 36 Cl deposition flux between the T1-T8 sites and the T10 site is 1.12 atoms m −2 s −1 , which is subtracted from the measurement of 11.6 ± 0.2 atoms m −2 s −1 at T10 to determine the gross 36 Cl deposition flux at T1-T8 sites (F 36gross ), achieving 10.5 atoms m −2 s −1 . The F 36gross is the 36 Cl dry deposition flux expected without chloride deficits, significantly higher than the measurements of 3.62-7.01 atoms m −2 s −1 at T1-T8 sites. There is a negative correlation between the measured 36 Cl deposition fluxes (F 36meas , atoms m −2 s −1 ) and %Cl deficit : F 36meas = −0.06 × %Cl deficit + 9.53 (n = 9, R 2 = 0.87, p < 0.01) (Figure 2a). Instead, no partial correlation (p > 0.05) existed between F 36meas and precipitation rates with chloride deficit effect removed, suggesting in regions where dry deposition is dominant, precipitation does not directly influence 36 Cl deposition fluxes. Thus, Equation 1 was adapted to account for the dependence on chloride deficits for the T1-T8 sites (≤10 mm yr −1 ): where S F (%Cl deficit ) is the slope (−0.06) of the F 36meas versus %Cl deficit . The F 36theor at T1-T8 sites were then obtained in the range of 5.18-8.70 atoms m −2 s −1 , generally in agreement with the measurements (Table 1). Next, we would briefly discuss possible local variables that might cause small deviations between F 36meas and F 36theor at each site.

The Coastal Site T1
The F 36meas of 6.64 ± 0.23 atoms m −2 s −1 at the T1 site is consistent with F 36theor of 6.62 atoms m −2 s −1 , though this site is heavily influenced by significant SSA deposition and fog deposition. Assuming that the chloride deposition at the T1 site was primarily from SSAs, the 36 Cl deposition flux contributed by the ocean was calculated to be 0.15 atoms m −2 s −1 by multiplying the global average 36 Cl/Cl ratio for oceanic chloride of 0.71(±0.08) × 10 −15 (Fifield et al., 2013) and the chloride deposition rate of 11.05 mmol m −2 yr −1 (Wang et al., 2014). This is only~2% of the measurement, suggesting trivial contribution of SSAs to 36 Cl deposition flux at the T1 site. Coastal fog of~22 L m −2 yr −1 could have contributed to 36 Cl deposition since fog can scavenge both aerosols (with chloride deficits) and atmospheric gases (including the displaced HCl (g) ), leading to no chloride deficits in local fog (Wang et al., 2014) as found in other studies (e.g., Kikuchi et al., 2009). However, there was a significant chloride deficit (60.5%, Table 1) in deposition at the T1 site, indicating the dominance of dry deposition of aerosols at T1. The large chloride deficit and relatively small 36 Cl deposition flux at the T1 site may be due to sizeable emissions of NO x and SO 2 from power/chemical plants and transport, associated with nearby cities of Mejillones and Antofagasta (Wang et al., 2014).

The Inland Sites T2-T6
The F 36meas at the T2-T6 sites (3.62(±0.18)-6.20(±0.21) atoms m −2 s −1 ) ( Table 1) were lower than F 36theor except at the T1, T2, and T4 sites (Table 1), likely due to the divergence in size fractions for 36 Cl and stable chlorine. While stable chlorine may preferably accumulate on coarse aerosols that originate from sea sprays or surface mineral dust (Li et al., 2019;Wang et al., 2014), 36 Cl mainly produced in the stratosphere tends to become associated with fine aerosols that are susceptible to downward transport during STT. This could explain while the chloride deposition rate decreased by an order of magnitude from the coastal T1 site to the inland T2-T6 sites (Table 1), the F 36meas at T2 (6.20 ± 0.21 atoms m −2 s −1 ) was only slightly lower than at the T1 site (6.64 ± 0.23 atoms m −2 s −1 ). However, fine aerosols were reported to have substantially larger chloride deficits (~90%) than coarse aerosols (~30%) due to larger surface to volume ratios (Harkel, 1997;Hsu et al., 2007), suggesting a confounding role of aerosol size distribution in affecting atmospheric 36 Cl inventory. Therefore, the overestimations of F 36meas based on Equation 2 (Table 1) at the T3, T5, and T6 sites might reflect more contributions of fine aerosols with larger chloride deficits, with uncertainties derived from corrections using bulk chloride deficits. Nevertheless, the lowest F 36meas at the T6 (3.62 ± 0.18 atoms m −2 s −1 ) and T3 (4.29 ± 0.14 atoms m −2 s −1 ) sites were associated with their largest chloride deficits (T6: 85.2%, T3: 81.7%), which were mainly accounted for by NO x emissions from Calama city and Chuquicamata smelter (~50 km to the north) near the T6 site and roads~2 km away from the T3 site (Li et al., 2019;Wang et al., 2014).

The Pre-Andean Sites T7-T8
The pre-Andean T7 or T8 sites had precipitation rates of~10 mm yr −1 during (DGA, 2010Houston, 2006). The T7 site had the second lowest F 36meas of 3.81 ± 0.13 atoms m −2 s −1 , which was associated with the still significant chloride deficit (68.1%, Table 1) and its lowest bulk dust and salt deposition rates among all the trap sites (Wang et al., 2014). The higher F 36meas of 7.01 ± 0.16 atoms m −2 s −1 at the T8 site was likely a result of lower chloride deficit (28.8%). However, the significantly lower F 36meas at the T7 and T8 sites than F 36theor of 6.25 and 8.70 atoms m −2 s −1 , respectively, could also be due to divergences in the chloride deficit between 36 Cl-laden fine aerosols and stable chlorine-rich coarse aerosols as stated above.

The Andean Site T10
The largest F 36meas (11.6 ± 0.2 atoms m −2 s −1 ) occurred at the T10 site, likely a consequence of strong STT and enhanced precipitation on the Andes as noted earlier. T10 did not have a chloride deficit, but a chloride surplus (i.e., chloride deficit: −30.0%), evidencing that not all chloride in deposition at the T10 site was associated with Na + , but some with Ca 2+ (Wang et al., 2014). The NO 3 − deposition rate at the T10 site was~4-9 times lower compared to at the other sites (Table 1), suggesting that NO x or HNO 3(g) did not effectively transport from the lowlands to this altitude (4,260 m), and therefore, F 36meas at T10 was least impacted by the chloride deficit.

Distribution of Meteoric 36 Cl/Cl Ratio in the Atacama
While variations in the 36 Cl deposition flux were attributed to regional STT, precipitation, and chloride deficits, variations in the meteoric 36 Cl/Cl ratio were mainly attributed to differences in local chloride sources. This is evidenced by the significant exponential correlation between meteoric 36 Cl/Cl ratios and chloride deposition rates at the T1-T10 sites (n=9, R 2 = 0.91, p < 0.01) (Figure 2b). The meteoric 36 Cl/Cl ratio at the T1 site (~6 km from the Pacific Ocean) of 31.5 ± 1.1 × 10 −15 was lowest, corresponding to the highest chloride deposition rate of 11.05 mmol m −2 yr −1 , which was attributed to the dilution by dry deposition of SSAs with a global average 36 Cl/Cl ratio of 0.71(±0.08) × 10 −15 (Fifield et al., 2013). The inland T2-T6 sites, except T5, had similar meteoric 36 Cl/Cl ratios ranging from 223(±11) × 10 −15 to 247(±10) × 10 −15 (Table 1).
Since there was no recorded precipitation at the T2-T6 sites during (DGA, 2010, these meteoric 36 Cl/Cl ratios could be interpreted as the 36 Cl/Cl ratios of dry deposition. The T5 site had a significantly lower meteoric 36 Cl/Cl ratio (146(±5) × 10 −15 ), which was attributed to its proximity (~7.5 km to the southwest) to Salar de Pampa Blanca (Stoertz & Ericksen, 1974;Wang et al., 2014) (Figure 1). The surface mineral in the crusts of Salar de Pampa Blanca had a low 36 Cl/Cl ratio of 18.0(±1.8) × 10 −15 (Table 1), close to the reported 36 Cl/Cl ratios for Atacamite minerals of 11.1 × 10 −15 -28.0 × 10 −15 (Reich et al., 2008), but lower than the secular equilibrium 36 Cl/Cl ratios of 53 × 10 −15 for vertical vein fill and 87 × 10 −15 for shallow groundwater in the same Baquedano region of the Atacama (Sturchio et al., 2009). The salar could act, in a fashion similar to the ocean, as a local source of chloride as mineral dust that is entrained and deposited nearby, lowering local 36 Cl/Cl ratios. However, dry deposition of 10-100 μm of particles (Seinfeld & Pandis, 2006) is dominated by gravitational settling; thus, reworking of surface chloride minerals from Salar de Pampa Blanca would mainly impact the nearby T5 rather than the other traps. Likewise, the meteoric 36 Cl/Cl ratios at the T7 (180 ± 6 × 10 −15 ) and T8 (94.0 ± 2.2 × 10 −15 ) sites were relatively low, likely a consequence of the proximities to two nearby open-pit mines south of the T7 site and the Salar de Atacama (Li et al., 2019). The Salar de Atacama, the region's largest halite-type salt playa dating back to the Oligocene-Miocene boundary (Stoertz & Ericksen, 1974), is to the southeast of the T7 (~35-km distance) and T8 (on the salar rim) sites. The surface minerals in the open-pit mines or the Salar de Atacama were not measured for 36 Cl/Cl ratios that, however, could be reasonably low. This is because the saline groundwaters accounting for the formation of surface chloride minerals in mines or salars are old (Corenthal et al., 2016), and once chloride entered the minerals at depths, the production of 36 Cl was limited, but decay did not cease (Reich et al., 2008). While the T10 site also had a relatively high chloride deposition rate (2.86 mmol m −2 yr −1 ), likely accounted for by the nearby Tara lake that had the Cl − content of~10 mmol L −1 and desiccated seasonally (Wang et al., 2014), its higher 36 Cl/Cl ratios than those at the T5, T7, and T8 sites were due to the large 36 Cl deposition flux facilitated by strong STT and enhanced precipitation.

Conclusion
This study constrained the meteoric 36 Cl/Cl ratios and 36 Cl deposition flux along a west-east transect across the Atacama Desert, Chile. The measured 36 Cl deposition flux ranged from 3.62 ± 0.18 to 11.6 ± 0.2 atoms m −2 s −1 , mainly from dry deposition, and are some of the lowest in the world, despite the influence of STT over the Andes. Variations in the 36 Cl deposition flux along the transect were mainly attributed to chloride deficits by acid displacement. The meteoric 36 Cl/Cl ratios varied eightfold across the transect, from 31.5 (±1.1) × 10 −15 to 247(±10) × 10 −15 , which was attributed to local inputs of chloride from the Pacific Ocean or surrounding salars. This study provides a baseline for the 36 Cl dry deposition in southern hemisphere tropical region and in hyperarid regions, where there are promising applications of meteoric 36 Cl as a potential tracer of local STT and duration of salt accumulation.

Data Availability Statement
All data are available in this manuscript or Wang et al. (2014).