Dissolved noble gas and isotopic tracers reveal vulnerability of groundwater in a small, high-elevation catchment to predicted climate changes



[1] Noble gas concentrations and multiple isotopic tracers in groundwater and stream water at a small, high-elevation catchment of the Sierra Nevada Mountains constrain recharge conditions and subsurface residence times of different groundwater components. We identify three sources that contribute to groundwater flow: (1) seasonal groundwater recharge with short travel times, (2) water with elevated radiogenic 4He that has experienced longer flow paths, and (3) upwelling of deep fluids that have “magmatic” helium and carbon isotope signatures. Results from our study illuminate two important aspects of the hydrological system that will have a direct impact on how this system responds to climate change: (1) recharge to the alluvial aquifer occurs primarily on the lower slopes of the catchment and is therefore sensitive to changes in snowline elevation and (2) deep groundwater in the western part of the aquifer is very young and provides very little buffering capacity. Although apparent groundwater ages indicate residence times range from less than a year to several decades, the water that recharges seasonally dominates the alluvial aquifer. Noble gas recharge temperatures are close to mean annual air temperature, and are 5°–11° higher than would be expected for direct influx of snowmelt. Excess air concentrations, indicating entrapment of air bubbles during recharge, are lower than would be expected for recharge through bedrock fractures. Instead, recharge likely occurs over vegetated areas on the lower slopes, as indicated by δ13C-dissolved inorganic carbon values that are consistent with incorporation of CO2 from soil respiration.

1. Introduction

[2] Predicted changes in the climate will have profound impacts on water resources and water management. Future climate changes in the western United States are likely to include a decrease in the percentage of precipitation that falls as snow; earlier onset of snowpack melting; an increase in number of rain on snow events; and changes in humidity, air temperature, and soil moisture [Dettinger and Cayan, 1995; Howat and Tulaczyk, 2005; Maurer and Duffy, 2005; Melack et al., 1997]. Snowmelt is an important component of groundwater recharge in high-elevation watersheds of the western United States [e.g., Earman et al., 2006]. In these watersheds, the predicted climate change impacts on snowmelt will likely alter the amount and timing of groundwater recharge, which may lead to reduced groundwater production, declining water tables, and reduced base flow to streams.

[3] Groundwater aquifers in alpine and subalpine basins play a critical role by storing and releasing snowmelt as base flow to streams long after seasonal precipitation and the disappearance of the snowpack and in this manner significantly impact streamflow and water temperature. Furthermore, geochemical hydrograph separations have shown that groundwater may supply a majority of alpine streamflow during peak snowmelt conditions [Liu et al., 2004]. Mountain-block aquifers can also provide significant recharge to mountain-front and basin-fill aquifers [Manning and Solomon, 2003; Manning and Solomon, 2005]. Despite being an important part of the water supply system, the recharge mechanisms, storage capacity, and residence times of high-elevation groundwater aquifers are poorly understood. The net change in recharge to mountain aquifers due to alterations in the timing of snowpack melting is not known in sign or magnitude, making it difficult to predict the response of these hydrological systems to climate change.

[4] Dissolved gas and isotope studies have given insights into the residence times and recharge processes operating in high-elevation watersheds. Manning and Caine [2007] used dissolved noble gas analyses to characterize groundwater recharge and residence times in a high-elevation (3300–3900 m above sea level [asl]) alpine watershed in Colorado. They determined that permeability decreases with depth and that aquifer parameters are relatively uniform throughout much of the watershed, with mean residence times between 8 and 11 years. Plummer et al. [2001] use dissolved gases to examine groundwater residence time in a mountainous region in Shenandoah National Park, VA. They found the shallow groundwater system to be dominated by young (<3 years) water and observed seasonally varying recharge temperatures, indicating shallow, seasonal recharge. In a study over a much larger geographic area, Manning and Solomon [2005] used dissolved noble gas results to examine mountain block recharge and subsurface flow to an adjacent basin. Rademacher et al. [2001] found that the apparent groundwater ages of springs in the Sagehen Basin, a small catchment approximately 27 km north of our study site, ranged from 1 to 36 years, and based on the chemical evolution of spring and creek waters, inferred that base flow to the local creek was dominated by moderately old groundwater [Rademacher et al., 2005]. Also in the Sagehen Basin, Blumhagen and Clark [2008] used carbon isotope compositions to show that dissolved inorganic carbon (DIC) in spring waters was inherited from respiration of CO2 in the soil zone.

[5] In alpine basins, the large gradients in altitude and temperature, which control gas solubility, make dissolved gases especially well suited to examining recharge processes and groundwater transport. Dissolved noble gases provide a snapshot of recharge water temperature and physical processes at the time of recharge and are transported conservatively in saturated media, providing a long-term record. Dissolved inorganic carbon concentrations, in combination with carbon isotope compositions, are useful for delineating the location of recharge and mixing of different water sources. When combined with measurements of tritium, helium isotopes provide a means of quantifying apparent groundwater subsurface residence time, or groundwater age, over a time scale relevant to the interaction between shallow groundwater and streamflow.

[6] In this paper the power of dissolved gas data to evaluate the vulnerability of water resources in high-elevation, snow-dominated watersheds is demonstrated in a small basin likely to experience altered runoff and recharge under warmer climate scenarios. Specifically, the questions that are addressed using dissolved gas and isotopic analyses are: What are the temperatures and time periods over which recharge takes place? Does water infiltrate through a soil layer or through fractures? What is the range in aquifer residence times for the bulk of the groundwater and for the groundwater most likely to contribute to stream base flow? One of the challenges in mountain hydrology is that many potential field sites lack access to adequate groundwater sampling points, such as monitoring wells. The research reported here takes advantage of the numerous monitoring wells and production wells in the Olympic Valley groundwater basin.

2. Study Site

[7] The Olympic Valley catchment is located 150 km east of Sacramento, CA, near Lake Tahoe in the Sierra Nevada, and has an area of approximately 22 km2 including alpine and subalpine zones. An alluvial aquifer extends eastward 4 km from the base of Granite Chief, a 2750 m peak that forms the center of Squaw Valley ski area, to the northward flowing Truckee River (Figure 1). The valley is drained by Squaw Creek, which is formed at the confluence of two major tributaries at the west margin of the basin at elevation 1898 m, entering the Truckee River at 1853 m. The groundwater basin is underlain by Cretaceous granites of the Sierra Nevada batholith, Jurassic metasediments, and Pliocene volcanics that also form the surrounding peaks. Glacial, lacustrine, and fluvial sediments fill the valley to a maximum thickness of 55 m near the center of the 0.8 km wide valley [Gasch and Associates, 1973]. A terminal moraine at the eastern end of the basin near the confluence of Squaw Creek and the Truckee River acted as a sediment dam throughout the Quaternary period.

Figure 1.

Topographic map of (a) the study area and (b) locations of the spring, wells, and faults discussed in the text. The outline of the alluvial aquifer is shown as a dotted line [after Hydrometrics-LLC, 2007]. Stream sampling sites as labeled in Table 1 are (1) Shirley Canyon, (2) South Fork, (3) Confluence, (4) Trapezoid, and (5) Squaw Creek Rd Bridge.

[8] The hydrogeology of the valley has been examined through drill core logs and surface exposures [Hydrometrics-LLC, 2007; Kleinfelder and Associates, 1987; West-Yost and Associates, 2003]. The unconsolidated valley-fill sediments act as an unconfined aquifer (dotted line on Figures 1a and 1b) except where laterally discontinuous fine-grained lacustrine deposits create semiconfined conditions. Three hydrostratigraphic units are loosely defined: a shallow unit consisting of fine-grained lake sediments and stream deposits, a middle unit of glacial sands and gravels, and a deep unit comprising fine-grained glacial lake sediments. Coarse-grained materials predominate in the western portion of the basin upstream of the production wells and are highly permeable. Sediments become less permeable in the downstream portion of the basin, and at the terminal moraine, groundwater occurrence is minimal. Estimates of hydraulic conductivity based on well tests range from 0.002 to 0.514 cm/s, averaging 0.067 cm/s [Kleinfelder Inc., 2000]. Groundwater also occurs in the crystalline rocks, with fractures providing secondary permeability. Four faults have been mapped across the valley based on surface exposures [Nevada Bureau of Mines and Geology, 2000], one of which is coincident with a spring (“Upwelling” spring on Figure 1) near monitoring wells 304/305.

[9] Precipitation in Olympic Valley occurs mainly in the form of snow in the winter months, with a smaller amount of precipitation occurring as rain during spring, summer, and fall. A U.S. Department of Agriculture SNOTEL site (site #784) located at 2447 m elevation in the catchment recorded an average annual precipitation of 1684 mm for the 1982–2008 water years (http://www.wcc.nrcs.usda.gov/snotel/). Estimated isohyetals [Di Luzio et al., 2008] show mean annual precipitation increasing from 1016 mm in the east to 1650 mm in higher elevations to the west.

[10] Olympic Valley is uniquely suited to a study of alpine and subalpine groundwater because of its relatively simple geometry and because many wells are available for sampling. However, the natural hydrologic cycle is altered through groundwater pumping and possibly because of channelization of the Squaw Creek streambed. Groundwater is extracted to supply the needs of valley residents and businesses (approximately 6.2 × 105 m3/yr), additional resorts, private residences and the ski area (extraction unknown), snow-making (8.5 × 104 m3/yr), and for irrigation of a golf course that fills a portion of the meadow surrounding Squaw Creek (2.5 × 105 m3/yr) [Hydrometrics-LLC, 2007]. The total average annual groundwater extraction by known sources was 8.9 × 105m3/yr for the years 1992–2004. Peak water demand occurs from July to October and is about twice the wintertime demand. The peak demand corresponds to the driest months of the year and may contribute to water table declines (Figure 2). The main production wells are located in a cluster in the western portion of the basin (Figure 1). Although the total groundwater discharge due to pumping is a small fraction of the annual precipitation that falls in the catchment (<3%), the cluster of production wells in the upstream portion of the alluvium does capture water that would contribute to down-gradient flow or to base flow in Squaw Creek. Fifteen monitoring well pairs are located on and around a golf course in the meadow that covers the lower valley. A small number of horizontal wells drilled into bedrock produce water at about 190 L/min, which is about 10% of the flow rate of the production wells in valley alluvium. Export of wastewater and increased evapotranspiration during irrigation cause a small net water loss to the catchment due to human activity. Total discharge from Squaw Creek is approximately 2 × 107 m3/yr [West-Yost and Associates, 2003], representing about 57% of the annual precipitation.

Figure 2.

(a) Daily snow water equivalent at SNOTEL Station #784 located at an elevation of 2447 m in the study area (http://www.wcc.nrcs.usda.gov/snotel/). (b) Maximum daily water level elevation during pumping at SVPSD Well 2 (http://www.svpsd.org/scada/aquiferwebdata.html). A dashed line shows the elevation of the Squaw Creek Bed where it passes near Well 2.

[11] Interaction between the groundwater production wells and streamflow is of concern because of the potential for adversely affecting habitat of brown trout and other fauna. The creek is gaining throughout the annual period of substantial streamflow (November–June), as indicated by water levels (Figure 2) and stream gauging stations monitored by SVPSD. During the summer months, when streamflow decreases to <0.05 m3/s, the groundwater elevation recorded in pumping production wells adjacent to the stream falls below the elevation of the creek bed (Figure 2). Hydrographs in production wells recorded during pumping since 1997 show total water level ranges of approximately 2.4m, with seasonal lows in late fall and highs in late spring [West-Yost and Associates, 2003]. Water level contours roughly mirror topography, with flows parallel to Squaw Creek in the center of the basin, and flow toward the creek along the margins of the basin. Head differences in paired monitoring wells show a general trend of slight downward gradients in the western portion of the meadow and upward gradients in the eastern portion [West-Yost and Associates, 2003]. During the year in which the study was carried out, the upper reach of the creek became dry by late summer, but deep pools and low flow persisted in the lower reach.

3. Methods

[12] This study includes results from eight production wells, including two horizontal wells located about 100 m above the valley floor, 22 monitoring wells, and 5 stream sampling sites (Figure 1 and Table 1). Samples for DIC were passed through a 0.45 μm filter and stored in 40 mL dark glass vials with no headspace. The DIC samples were kept cold in the field and stored in a refrigerator until analysis. Tritium samples were collected in 1 L glass containers with plastic caps. Dissolved noble gas samples (He, Ne, Ar, Kr, Xe) were collected using clear Tygon tubing to connect the sample vessel (8 mm inner diameter copper tubing, 250 mm long) to the wellhead of operating production wells or monitoring wells pumped by a Grundfos® submersible pump. Horizontal wells drilled into bedrock were likewise sampled at the wellhead. Water flowed for several minutes to purge air from the sample tube. The copper tubing was tapped lightly to dislodge bubbles and a visual inspection for bubbles was made. Close attention was paid to maintaining sufficient pressure in the sampling apparatus, and backpressure was applied when necessary to prevent escape of dissolved gas. Steel clamps pinched the copper tubing flat in two locations to secure the water sample.

Table 1. Measurements of Dissolved Gas and Isotopic Compositions From Horizontal Wells (HW), Monitoring Wells (MW), Production Wells (PW), a Spring (SP), and Stream Waters (SW)a
Sample SiteCollection DateSample TypeScreen Depth (m)Approximate Bedrock Depth (m)DIC (mg/L C)δ13C (‰Vienna Peedee Belemnite)3H (pCi/L)3He/4He (10−6)4He (10−8 cm3 STP/g)Ar (10−4 cm3 STP/g)Kr (10−8 cm3 STP/g)Ne (10−8 cm3 STP/g)Xe (10−8 cm3 STP/g)
  • a

    NM, Measurement failed.

Horizontal Well 106/19/2008HW  26−17.511.9 ± 0.51.64 ± 0.056.49 ± 0.214.00 ± 0.088.99 ± 0.2728.33 ± 0.571.25 ± 0.04
Horizontal Well 206/19/2008HW  32−17.619.8 ± 1.21.77 ± 0.066.07 ± 0.194.15 ± 0.089.22 ± 0.2827.64 ± 0.551.28 ± 0.04
MW 30105/13/2008MW3–53438−15.78.4 ± 0.61.65 ± 0.013.56 ± 0.073.42 ± 0.077.75 ± 0.2319.78 ± 0.401.15 ± 0.03
MW 30205/13/2008MW12–143440−14.98.3 ± 0.42.09 ± 0.026.79 ± 0.143.39 ± 0.077.66 ± 0.2321.11 ± 0.421.08 ± 0.03
MW 30306/18/2008MW3–53020−18.810.5 ± 1.31.46 ± 0.046.67 ± 0.133.53 ± 0.077.78 ± 0.2324.93 ± 0.501.12 ± 0.03
MW 30405/14/2008MW3–51273−7.76.1 ± 0.62.19 ± 0.0295.29 ± 1.913.85 ± 0.088.39 ± 0.2528.02 ± 0.561.24 ± 0.04
MW 30505/14/2008MW11–131272−8.16.2 ± 0.62.15 ± 0.02101.83 ± 2.043.87 ± 0.088.53 ± 0.2626.42 ± 0.531.20 ± 0.04
MW 30605/14/2008MW2–32321−20.810.3 ± 1.21.52 ± 0.015.74 ± 0.114.11 ± 0.089.09 ± 0.2729.85 ± 0.601.31 ± 0.04
MW 30906/18/2008MW3–52422−14.416.2 ± 1.3      
MW 31006/18/2008MW11–122422−20.99.6 ± 1.21.64 ± 0.055.30 ± 0.113.75 ± 0.088.19 ± 0.2526.67 ± 0.531.18 ± 0.04
MW 31506/18/2008MW2–324774.211.1 ± 1.0      
MW 31606/18/2008MW8–102466−16.69.5 ± 1.3      
MW 31805/13/2008MW3–54167−0.716.2 ± 1.1      
MW 32105/14/2008MW3–53429−17.59.8 ± 1.21.68 ± 0.015.28 ± 0.113.33 ± 0.077.61 ± 0.2324.06 ± 0.481.08 ± 0.03
MW 32205/14/2008MW17–183421−18.113.2 ± 1.31.65 ± 0.017.06 ± 0.143.82 ± 0.088.58 ± 0.2626.64 ± 0.531.24 ± 0.04
MW 32705/14/2008MW11–131583−8.1       
MW 32805/14/2008MW3–51523−18.311.3 ± 1.31.42 ± 0.015.37 ± 0.113.68 ± 0.078.67 ± 0.2629.56 ± 0.591.25 ± 0.04
MW 32905/14/2008MW13–1523203−5.70.3 ± 0.62.77 ± 0.03290.88 ± 5.824.26 ± 0.098.95 ± 0.2731.69 ± 0.631.32 ± 0.04
MW 33005/13/2008MW8–932222−4.7−0.5 ± 0.8      
MW 33105/13/2008MW2–33220−17.111.5 ± 0.91.51 ± 0.016.19 ± 0.123.50 ± 0.078.02 ± 0.2421.29 ± 0.431.14 ± 0.03
MW 5D05/13/2008MW12–143418−19.38.4 ± 0.61.36 ± 0.017.38 ± 0.153.63 ± 0.077.95 ± 0.2425.76 ± 0.521.14 ± 0.03
MW 5D08/12/2008MW12–143416−18.89.0 ± 0.91.39 ± 0.086.74 ± 0.553.53 ± 0.078.03 ± 0.2425.16 ± 0.501.12 ± 0.03
MW 5S05/13/2008MW6–83428−19.611.0 ± 0.71.37 ± 0.014.84 ± 0.103.81 ± 0.088.58 ± 0.2622.91 ± 0.461.16 ± 0.03
MW 5S08/12/2008MW6–83425−19.211.6 ± 1.01.37 ± 0.084.54 ± 0.373.20 ± 0.067.39 ± 0.2220.77 ± 0.421.03 ± 0.03
MW-PJOW05/13/2008MW24–26?11−13.09.8 ± 0.91.38 ± 0.014.74 ± 0.093.43 ± 0.078.47 ± 0.2518.67 ± 0.371.24 ± 0.04
MW-PJOW08/12/2008MW24–26?11−12.99.2 ± 0.91.36 ± 0.084.59 ± 0.373.48 ± 0.078.65 ± 0.2620.08 ± 0.401.26 ± 0.04
MW T409/26/2008MW16–484422−12.58.3 ± 0.51.54 ± 0.0215.67 ± 0.313.72 ± 0.078.36 ± 0.2526.36 ± 0.531.15 ± 0.04
MWC Well 104/25/2008PW18–303822−18.49.6 ± 0.41.37 ± 0.01NM3.69 ± 0.078.22 ± 0.2524.84 ± 0.501.14 ± 0.03
MWC Well 108/12/2008PW18–303823−19.111.0 ± 0.91.43 ± 0.085.67 ± 0.463.57 ± 0.078.12 ± 0.2423.68 ± 0.471.16 ± 0.03
MWC Well 204/25/2008PW11–303830−19.09.7 ± 0.41.36 ± 0.01NM3.83 ± 0.088.22 ± 0.2523.06 ± 0.461.23 ± 0.04
MWC Well 208/12/2008PW11–303825−19.011.5 ± 0.91.37 ± 0.045.34 ± 0.173.64 ± 0.078.30 ± 0.2524.06 ± 0.481.14 ± 0.03
SVPSD Well104/25/2008PW23–343721−16.99.3 ± 0.4      
SVPSD Well106/18/2008PW23–343721−16.59.1 ± 2.41.41 ± 0.047.50 ± 0.153.64 ± 0.078.29 ± 0.2522.49 ± 0.451.19 ± 0.04
SVPSD Well108/12/2008PW23–343720−16.48.3 ± 0.81.41 ± 0.087.99 ± 0.653.75 ± 0.078.96 ± 0.2723.99 ± 0.481.24 ± 0.04
SVPSD Well204/25/2008PW10–232417−17.210.9 ± 0.5      
SVPSD Well206/18/2008PW10–232417−16.79.9 ± 0.91.44 ± 0.045.77 ± 0.123.65 ± 0.078.49 ± 0.2522.59 ± 0.451.23 ± 0.04
SVPSD Well208/12/2008PW10–232416−18.110.4 ± 0.51.41 ± 0.085.56 ± 0.453.50 ± 0.078.49 ± 0.2521.15 ± 0.421.19 ± 0.04
SVPSD Well304/25/2008PW22–354617−17.410.0 ± 0.41.35 ± 0.01NM3.65 ± 0.077.95 ± 0.2421.47 ± 0.431.18 ± 0.04
SVPSD Well306/18/2008PW22–354618−16.910.3 ± 1.01.41 ± 0.045.43 ± 0.113.56 ± 0.078.29 ± 0.2521.76 ± 0.441.20 ± 0.04
SVPSD Well504/25/2008PW22–394617−18.39.9 ± 0.41.36 ± 0.015.75 ± 0.113.63 ± 0.077.97 ± 0.2422.78 ± 0.461.19 ± 0.04
SVPSD Well506/18/2008PW22–394617−17.914.3 ± 2.41.39 ± 0.045.88 ± 0.123.74 ± 0.078.21 ± 0.2525.82 ± 0.521.16 ± 0.03
SVPSD Well508/12/2008PW22–394616−17.312.6 ± 1.01.40 ± 0.045.49 ± 0.183.49 ± 0.077.99 ± 0.2422.04 ± 0.441.17 ± 0.04
Upwelling04/25/2008SP  33–16.5       
Upwelling05/14/2008SP  33−15.610.2 ± 0.9      
Upwelling08/12/2008SP  26−18.0       
Confluence04/25/2008SW  7−10.29.5 ± 0.4      
Confluence05/14/2008SW  4−4.811.2 ± 0.7      
Confluence06/19/2008SW  4−3.4       
Shirley Canyon04/25/2008SW  5−5.39.5 ± 0.4      
Shirley Canyon05/14/2008SW  4−4.811.5 ± 0.6      
Shirley Canyon06/18/2008SW  3−2.7       
South Fork at Bridge04/25/2008SW  6−7.511.2 ± 0.5      
South Fork at Bridge05/14/2008SW  4−6.0       
South Fork at Bridge06/18/2008SW  5−2.6       
Squaw Creek Rd bridge04/25/2008SW  8−10.211.0 ± 0.5      
Squaw Creek Rd bridge05/14/2008SW  4−7.3       
Squaw Creek Rd bridge06/19/2008SW  5−5.9       
Squaw Creek Rd bridge08/12/2008SW  11−11.0       
Squaw Creek Rd bridge09/25/2008SW  13−10.19.2 ± 0.5      
Trapezoid04/25/2008SW  7−8.811.5 ± 0.5      

[13] All analyses were performed at Lawrence Livermore National Laboratory (LLNL). Dissolved inorganic carbon and its carbon isotope composition were determined using the automated DIC-dissolved organic carbon-isotope ratio mass spectrometry technique [St-Jean, 2003] consisting of an OI Analytical Model 1030 Carbon analyzer and a Micromass (now Isoprime Ltd) IsoPrime isotope ratio mass spectrometer. Carbon isotope compositions (13C/12C) are reported as delta values in per mil relative to the Vienna Peedee Belemnite reference, with an analytical uncertainty of ±0.3‰. Copper tube samples for noble gas analysis were mounted on a multiport gas handling manifold under vacuum. Reactive gases were removed with multiple reactive metal getters. Known quantities of isotopically enriched 22Ne, 86Kr, and 136Xe were added to provide internal standards. The isotope dilution protocol used for measuring noble gas concentrations is insensitive to potential isotopic composition variation in dissolved gases (especially Ne) due to diffusive gas exchange. Noble gases were separated from one another using cryogenic adsorption. Helium was analyzed using a VG-5400 noble gas mass spectrometer. Other noble gas isotopic compositions were measured using a quadrupole mass spectrometer. The Ar abundance was determined by measuring the total noble gas sample pressure using a high-sensitivity capacitance manometer. The procedure was calibrated using water samples equilibrated with the atmosphere at a known temperature and air standards spiked with known quantities of the noble gases. Tritium concentrations were determined on 500 g subsamples by the 3He in-growth method (approximately 25 day accumulation time). Analytical uncertainties are approximately 1% for 3He/4He; 2% for He, Ne, and Ar; and 3% for Kr and Xe. Errors for derived parameters such as groundwater age and recharge temperature are propagated using analytical errors for the individual measured quantities (Table 2). A detailed description of the data reduction routine is reported in the report by Ekwurzel [2004].

Table 2. Calculations Based on Noble Gas Concentrations and Isotope Ratiosa
Well NameDateScreen Depth (m)ΔNe4Herad (10−9 cm3 STP/g)NGRTb (°C)χ2Agec% Pre-modernGroup
  • a

    For date, read 5/13/08 as 13 May 2008.

  • b

    An elevation of 1950 m used to determine pressure, except at the two horizontal wells, where an elevation of 2050 m is applied.

  • c

    A crustal 3He/4He ratio of 6 × 10−7 is used in the age determination for samples with [Hemeas − (Hesol + Hexs air)] >2 × 10−9 cm3 STP/g.

Upper Valley MWs
MW-PJOW5/13/0824–2617%4.5 ± 1.65.5 ± 0.61.01 ± 1<20%1
MW-PJOW8/12/0824–2624%<25.3 ± 0.63.1<1 ± 1<20%1
MW 5D5/13/0812–1438%11.8 ± 4.19.2 ± 0.90.315 ± 552%2
MW 5D8/12/0812–1437%6.3 ± 3.89.5 ± 0.90.811 ± 936%2
MW 5S5/13/086–830%<28.2 ± 0.711.0<1 ± 1<10%1
MW 5S8/12/086–825%<211.4 ± 0.80.5<1 ± 1<10%1
MW T49/26/0816–4837%94.0 ± 4.88.8 ± 0.90.449 ± 398%2
Production Wells
SVPSD Well16/18/0823–3427%22.3 ± 2.87.2 ± 0.70.423 ± 475%2
SVPSD Well18/12/0823–3431%23.2 ± 3.46.2 ± 0.71.223 ± 673%2
SVPSD Well26/18/0810–2327%4.8 ± 2.86.4 ± 0.70.011 ± 421%2
SVPSD Well28/12/0810–2323%5.8 ± 2.37.1 ± 0.71.311 ± 631%2
SVPSD Well34/25/0822–3523%NM7.3 ± 0.73.0<1 ± 1<10%1
SVPSD Well36/18/0822–3525%3.2 ± 2.56.8 ± 0.70.14 ± 216%1
SVPSD Well54/25/0822–3928%3.8 ± 2.97.3 ± 0.71.3<1 ± 1<10%1
SVPSD Well56/18/0822–3937%<28.6 ± 0.80.51 ± 1<10%1
SVPSD Well58/12/0822–3927%2.7 ± 2.67.7 ± 0.70.83 ± 1<10%1
MWC Well 14/25/0818–3027%<28.6 ± 0.84.1<1 ± 1<10%1
MWC Well 18/12/0818–3032%<28.1 ± 0.80.26 ± 6<10%1
MWC Well 24/25/0811–3022%<26.2 ± 0.77.5<1 ± 1<10%1
MWC Well 28/12/0811–3031%<28.7 ± 0.81.4<1 ± 1<10%1
Horizontal Wells
Horizontal Well 16/19/08 41%<26.5 ± 0.80.615 ± 2<10%1
Horizontal Well 26/19/08 40%<25.6 ± 0.83.517 ± 2<10%1
Shallow Lower Valley MWs (2–5 m BGS)
MW 3015/13/083–516%<27.7 ± 0.70.717 ± 160%1
MW 3036/18/083–537%6.2 ± 3.79.5 ± 0.81.010 ± 245%2
MW 3045/14/083–543%885 ± 187.0 ± 0.92.03He excessNA3
MW 3065/14/082–343%<25.6 ± 0.80.113 ± 135%1
MW 3215/14/083–532%<210.3 ± 0.82.621 ± 267%1
MW 3285/14/083–544%<2  4 ± 1<10%1
MW 3315/13/082–324%11.8 ± 2.38.2 ± 0.70.412 ± 158%2
Deep Lower Valley MWs (1118 m BGS)
MW 3025/13/0812–1424%18.0 ± 2.39.7 ± 0.70.438 ± 197%2
MW 3055/14/0811–1339%954 ± 207.7 ± 0.80.43He excessNA3
MW 3106/18/0811–1240%<28.2 ± 0.90.519 ± 362%1
MW 3225/14/0817–1838%7.3 ± 4.56.7 ± 0.80.319 ± 161%2
MW 3275/14/0811–1375%803 ± 259.2 ± 2.50.13He excessNA3
MW 3295/14/0813–1549%2832 ± 575.9 ± 0.90.83He excessNA3

4. Results and Discussion

[14] A total of 34 samples from 25 wells were analyzed for noble gas, and 50 samples from wells and surface water were analyzed for tritium (Table 1). Samples were collected between April and September of 2008. A total of 60 DIC samples were collected from 30 wells, 1 spring, and 5 stream sampling sites (Table 1).

4.1. Excess Air

[15] The concentration of dissolved noble gases in groundwater is virtually always greater than equilibrium solubility. The portion of gas in excess of equilibrium solubility is termed “excess air” because of its compositional similarity to air [Aeschbach-Hertig et al., 2000; Holocher et al., 2002]. During transport through the unsaturated zone, infiltrating water may entrain or trap air bubbles that subsequently dissolve in groundwater. Air bubbles may also become trapped in groundwater during fluctuations in water table elevation. The concentration of excess air provides unique information about recharge processes, including the degree to which infiltrating water incorporates unsaturated zone gas. For dissolved noble gases, addition of excess air has the greatest relative impact on He and Ne concentrations because the equilibrium solubility components of these gases are relatively small.

[16] The total measured concentration of dissolved gas i, Ci,m, is the sum of multiple components [Lehmann et al., 1993]:

equation image

where subscripts exc, rad, and ter refer to excess air, radiogenic, and terrigenic components, respectively. The equilibrium component for each gas is determined from Henry's law. Radiogenic and terrigenic contributions to Ne, Ar, Kr, and Xe are assumed to be negligible in the study area.

[17] A common way to represent the amount of excess air is as percent excess Ne, or ΔNe (excess Ne relative to equilibrium component; Table 2). Neon concentrations are used in determining excess air because Ne can be assumed to derive solely from the atmosphere and because Ne is measured with high precision. Excess air may be fractionated during the recharge process (whereby lighter gases are depleted relative to heavier gases) and Aeschbach-Hertig et al. [1999], and later Cey et al. [2008], examined optimization models to treat fractionated excess air and calculate noble gas recharge temperatures. For samples from Olympic Valley, gas concentration ratios (e.g., He/Ne versus Ar/Ne) indicate that addition of excess air above equilibrium solubility is with gas of very nearly atmospheric composition, and therefore, the unfractionated excess air model was used to evaluate the noble gas concentration results. Use of one of the other excess air models would result in changes to calculated groundwater ages and recharge temperatures of not more than +2 years and +1°C, respectively.

[18] The total concentration of gas i as given by the unfractionated air model is [Heaton and Vogel, 1981]:

equation image

where Ad is concentration of dry air dissolved, and zi is the volume fraction of gas i in dry air (volume fraction composition of dry air from Ozima and Podosek [1983]). Additional equations and solubility coefficients used to calculate equilibrium solubility and excess air components are given in the report by Ekwurzel [2004].

[19] Recently, Cey et al. [2008] interpreted over 900 analyses of excess air in groundwater samples from the major groundwater basins in California. Results for over 400 samples not affected by artificial recharge (which can greatly increase excess air due to rapid changes in hydraulic head and water table height) are shown in Figure 3, compared with excess air results for 35 samples from Olympic Valley and 31 wells from Handcart Gulch, an alpine watershed in Colorado [Manning and Caine, 2007]. Excess air in alpine basins is relatively unexplored, but Manning and Caine [2007] find high excess air concentrations, as shown in Figure 3, associated with bedrock wells, and attribute the high values to recharge through fractures. Similarly, Plummer et al. [2001] found higher concentrations of excess air in groundwater from wells in fractured rock compared to other wells in Shenandoah National Park. The highly dynamic fluctuations in water table elevation during recharge through fractured rock lead to elevated excess air concentrations [Ingram et al., 2007]. Very low excess air concentrations are expected where an unsaturated zone is not present, as for continuous stream recharge where the water table intersects the surface [Beyerle et al., 1999; Kipfer et al., 2002]. If an unsaturated zone is present between the stream and the water table, somewhat higher excess air concentrations are expected.

Figure 3.

Observed ranges in excess air, expressed as ΔNe, for samples from Olympic Valley (categorized by well type), samples from the alpine watershed Handcart Gulch, and a large number of samples from around California sampled under the GAMA program. Handcart Gulch after Manning and Caine [2007]; California GAMA after Cey et al. [2008].

[20] Excess air concentrations in Olympic Valley wells range from ΔNe of 16%–75%, with an average value of 34% and a median value of 31%, in line with the median value of 28% for the large California data set. These values argue against substantial recharge through fractures. The observed range for Olympic Valley samples suggests that groundwater recharges through an unsaturated zone but not under conditions of high hydraulic head, large fluctuations in water table height, or very high infiltration rates. Interestingly, this observation holds true for horizontal well samples, which are drilled into fractured bedrock, indicating that recharge through soils overlying the bedrock is likely even at higher elevation in this basin. The moderate average excess air concentrations observed in Olympic Valley wells also argue against stream recharge as the predominant recharge mechanism, since lower excess air values would be expected where there was a direct connection between the stream and the water table. Water table elevations in the western part of the aquifer are typically higher than the elevation of Squaw Creek during periods of creek flow (Figure 2), so recharge through an unsaturated zone under the creek is unlikely.

4.2. Recharge Temperature

[21] Solubilities of the noble gases in fresh water vary as a function of temperature and pressure according to Henry's law and are well-known from theoretical and empirical studies [Andrews, 1992]. The strong temperature dependence, especially for the heavy gases, makes it possible to determine the temperature at the water table at the time of recharge [Stute and Schlosser, 1999]. Cey et al. [2009] demonstrate that recharge temperatures determined from noble gas concentrations are very close to measured water table temperatures. Using inverse modeling, the recharge temperature and pressure are determined from noble gas concentrations. However, simultaneous estimation of both temperature and pressure results in a poorly resolved solution because these two parameters are strongly correlated. In general, pressure (or elevation) is more easily constrained by geographic conditions; once an elevation is estimated, temperature is well constrained. In wells producing water of mixed age, noble gas recharge temperatures (NGRTs) represent mean, integrated values for the mixtures.

[22] NGRTs are calculated from Xe concentrations, after subtraction of the excess air component, using a polynomial fit to the Xe-temperature solubility curve. Xenon is used to calculate NGRTs because Xe solubility has the strongest temperature dependence. Equilibrium concentrations of Ar and Kr are then calculated using a polynomial that gives a concentration based on the recharge temperature determined from Xe, multiplied by a pressure fraction, which is based on the assumed recharge elevation.

[23] The calculated Ar and Kr values are compared with the measured values and used to check the goodness of fit between measured and modeled values, which is reported as χ2 (Table 2). This situation has 2 d.f. so a χ2 value >6 corresponds to a probability (p) value of <0.05. The p value is the probability that the deviation between modeled recharge temperatures and their true values is solely due to the 1 sigma measurement error. The data reduction procedure for calculating NGRTs and examining the goodness of fit is given in the report by Ekwurzel [2004].

[24] For this sample set, χ2 values have a mean of 1.54, and for 27 out of 35 samples, χ2 values are <2 (Table 2); thus, the model for recharge temperature calculation describes these data adequately. Four samples with somewhat higher χ2 values are included, but their calculated NGRTs have higher associated uncertainties. NGRTs, calculated assuming a recharge elevation of 1950 m, range from 5.3°C ± 0.6°C to 11.4°C ± 0.8°C, with an average value of 7.7°C. The assumed recharge elevation of 1950 m is close to the break in slope between the surrounding mountainous area and the valley floor, and about 50 m above wellhead elevations, which is a likely elevation of recharge for the basin groundwaters.

[25] Increasing the elevation in the recharge calculations will lead to lower calculated NGRTs. We can therefore calculate a minimum recharge temperature by assuming that the sample was recharged at the very top of the catchment (2750 m asl). Assuming the maximum recharge elevation for the catchment results in a minimum recharge temperature of 3.0°C ± 0.6°C. The highest recharge temperature, assuming maximum elevation, is 8.9°C ± 0.7°C, and the mean of all the samples is lowered to 5.4°C. It is unlikely that any significant proportion of the groundwater is recharged at this high elevation given the very small area available. A wider range of recharge elevations that covers much of the surface area, from 2300 to 1950 m, gives a mean recharge temperature that ranges from 6.7°C to 7.7°C. Given that this difference is close to the analytical uncertainty for most measurements, we use 1950 m as an estimate for elevation when calculating recharge temperatures for most wells. We use a recharge elevation of 2050 m for the horizontal wells, since they are located approximately 100 m above the valley floor.

[26] Even the minimum possible recharge temperatures calculated for the Olympic Valley wells are significantly higher than the melting point of snow and ice. If we assume that most recharge originates as snowmelt, then the water temperature must have increased prior to reaching the water table. This observation would indicate that an unsaturated zone is present during recharge rather than a direct connection between groundwater and the land surface and that the infiltrating water has a residence time that allows for equilibration with the shallow ground temperature.

[27] In general, soil temperatures near the surface show a damped version of surface temperature variations, but deeper in the unsaturated zone, temperatures approach the mean annual air temperature (MAAT) [e.g., Cey et al., 2009; Kipfer et al., 2002]. Flint et al. [2008] measured soil temperatures at a site near Yosemite at a similar elevation (2130 m asl) to the Olympic Valley and found that soil temperatures were fairly stable and increased with depth from approximately 1°C at 10 cm to approximately 3°C at 72 cm under a melting snowpack. Once the snowpack disappeared, Flint et al. [2008] observed rapid increases in soil temperatures of approximately 8°C–23°C, with the 10 cm depth showing diurnal temperature fluctuations.

[28] The recharge temperatures calculated for Olympic Valley wells fall close to or slightly above the long-term MAATs reported for the nearby Tahoe City NASA GISS Climate Station (http://data.giss.nasa.gov/gistemp/station_data/), which is located approximately 7 km to the southeast of Olympic Valley at an elevation of 1899 m (Figure 4). In most cases shallow ground temperatures are slightly greater (1°C–2°C) than the MAAT, so it is common for NGRTs to be slightly greater than MAAT [Kipfer et al., 2002]. The average maximum recharge temperature for groundwater samples (7.7°C) falls very close to the long-term average air temperature for May (7.8°C), when there is significant snowmelt (Figure 4c). The higher recharge temperatures fall closer to the long-term mean air temperature in June (11.4°C). This overlap between recharge temperatures and monthly mean air temperatures is consistent with recharge taking place over a 2–3 month period during the snowmelt season. However, a seasonal signal in the recharge temperatures would only be recorded if water table depths in recharge areas were shallow enough to be influenced by seasonal temperature. It is equally possible that the recharge temperatures primarily reflect MAAT, with some locally higher temperatures due to shallow water table depths. There is no discernable trend toward higher or lower recharge temperature values with apparent age (Figure 4b). However, the large range in NGRTs observed in samples with mean apparent ages ≤1 year (discussed in the next section) likely shows the effects of seasonal recharge, with lower temperatures reflecting recharge at the beginning of the melting period and higher temperatures reflecting late season recharge. In wells sampled more than once, only two wells show a significant seasonal change in recharge temperature (Table 2). The shallow well directly adjacent to the creek (Well 5S), changed from 8.2°C ± 0.7°C on 13 May to 11.4°C ± 0.8°C on 12 August, and the MWC Well 2 production well changed from 6.2°C ± 0.7°C on 25 April to 8.7°C ± 0.8°C on 12 August.

Figure 4.

Noble gas recharge temperatures and apparent groundwater ages shown with (a) mean annual air temperature, (b) quarterly air temperature, and (c) average monthly air temperature at the NASA/Goddard Institute for Space Studies station at Tahoe City, CA. The range (shaded) and average (horizontal line) of recharge temperatures, along with the typical snowmelt season are shown on Figure 4c.

4.3. Tritium Concentrations and Groundwater Ages

[29] All but two of the samples have tritium (3H; half-life 12.32 years) concentrations above the detection limit of approximately 1 pCi/L (Table 1), signaling the presence of groundwater recharged within the last 50 years. Most of the tritium concentrations overlap with the expected range for modern day precipitation, which limits the utility of tritium concentrations alone for determining ages. This range also overlaps with the tritium concentrations measured in Squaw Creek, which ranged from 9.2 to 11.5 pCi/L, with a mean for eight measurements of 10.6 pCi/L (Table 1). Well values at the high end of the observed range are the result of a contribution from global fallout from nuclear testing, while observed low values are the result of decay, and/or dilution with older, tritium-free water.

[30] Mean apparent groundwater ages, calculated from tritium and tritiogenic 3He concentrations, are shown in Table 2. In order to determine tritiogenic 3He, the measured 3He and 4He must be adjusted for contributions from the atmosphere (equilibrium solubility and excess air) and from subsurface sources [Cook and Solomon, 1997; Ekwurzel et al., 1994; Schlosser et al., 1989; Schlosser et al., 1988]. A significant buildup of radiogenic 4He due to decay of U and Th in crustal rocks takes place as the saturated zone residence time increases [Andrews et al., 1985; Torgersen and Clarke, 1985]. Radiogenic 4He is the portion of measured 4He remaining after subtracting solubility and excess air components. In addition, magmatic fluids can contribute dissolved helium that has a much higher 3He/4He ratio than atmospheric or crustal sources. The observed tritium and dissolved helium compositions indicate that all of these components are present and will be considered in the analysis of Olympic Valley groundwater.

[31] Wells with long screens typically sample water of differing ages. In examining samples with a mixture of ages, it is useful to determine the fraction of the mixture that recharged before about 1950 (the time of large increases in global atmospheric 3H due to nuclear weapons testing). The reported tritium-helium age is the mean apparent age of the portion of the sample that contains tritium above the detection limit. A rough estimate of the “percent pre-modern” is determined by comparing the initial 3H in a sample (i.e., measured 3H + 3Hetritiogenic) with the 3H in precipitation at the time and location of recharge (Figure 5). The nearest International Atomic Energy Agency GNIP stations where 3H in precipitation data were collected are Santa Maria, CA, and Menlo Park, CA (http://www-naweb.iaea.org/napc/ih/GNIP/IHS_GNIP.html), but these are incomplete records in very different physiographic settings than Olympic Valley. Figure 5 shows an exponential fit to mean annual averaged data for Western North America, along with measured values from two cities with long records. Several points for groundwater samples from Olympic Valley with apparent ages older than 10 years fall below the curve. The percentage of pre-modern water (Table 2) is calculated according to the difference between the expected value on the curve, and the observed value below the curve. This approach assumes that the sample is a binary mixture between pre-modern water and modern water of a single age and that no mixing of groundwater with different ages occurs during transport (i.e., “piston flow”). For exponential mixing of post-bomb pulse peak water [Cook and Böhlke, 1999], the initial 3H curve is generally somewhat higher than the smoothed curve, so larger pre-modern percentages would be calculated. In any case, samples with the oldest tritium-helium ages are mixed with a significant component of pre-modern water.

Figure 5.

Tritium concentrations measured in precipitation at two locations where long International Atomic Energy Agency GNIP records exist, along with an exponential curve that approximates mean annual values from western North America. Results for Olympic Valley well samples are plotted according to the calculated apparent tritium helium age (recharge year) and the measured tritium + tritiogenic 3He (initial 3H). Points that fall well below the curve contain a significant component of pre-modern water.

[32] Olympic Valley groundwater samples fall into three groupings: (1) samples with very young (≤1 year) to young (1–17 years) mean apparent ages and little pre-modern water; (2) samples with a component of relatively young (<50 years), tritiated water mixed with older, pre-modern water containing radiogenic 4He; and (3) samples with a smaller component of young water, mixed with older water and affected by gasses from a magmatic source.

[33] The three groups are distinguishable on a plot of 3He/4He versus Ne/He (Figure 6). On this plot, Ne and He concentrations are adjusted by subtracting the excess air component of each. For comparison, values for air saturated water at 8°C are also shown. Ne has only an atmospheric source, whereas He may be affected by the buildup of crustal He, accumulation of tritiogenic 3He, or addition of magmatic He. Crustal He can contribute both 3He (via an α,n reaction on 6Li) and 4He (via α decay of natural U and Th); the effect is insignificant for 3He but can be very large for 4He, so an increase in crustal He results in a decrease in Ne/He. Magmatic He sources have high 3He/4He and typically high He concentrations. On Figure 6, samples with magmatic He stand out as having high 3He/4He and low Ne/He, whereas samples with a component of radiogenic 4He have Ne/He ratios lower than air saturated water (corresponding to wells with radiogenic 4He concentrations >2 × 10−9 cm3 STP/g H2O in Table 2). Addition of tritiogenic 3He due to decay of 3H causes an increase in 3He/4He above values expected for atmospheric sources of He.

Figure 6.

A plot of the ratio of Ne/He (measured concentration minus amount due to excess air) versus the measured 3He/4He ratio. Samples in Group 1 with very young ages are close to solubility values; significant amounts of tritiogenic 3He bring some samples above the solubility ratio of 1.364 × 10−6. Group 2 samples are affected by crustal He, which results in a decrease in Ne/He and by tritiogenic 3He. Group 3 samples are affected by magmatic He. Orange triangle represents air saturated water (ASW) at 8°C.

4.3.1. Recent Recharge

[34] Group 1 (Figure 6) includes two monitoring wells and four production wells that yield water with a mean apparent age of less than 1 year (with a 2 sigma analytical uncertainty of about 1 year) and no detectable pre-modern water. Groundwater ages of 1 year or less in long screened, high-flow wells are unusual; fewer than 2% of drinking water wells (n = 1317) examined under California's Groundwater Ambient Monitoring and Assessment (GAMA) program (http://www.swrcb.ca.gov/gama/) have apparent tritium-helium ages ≤1 year. The wells producing very recent recharge are located in the western portion of the basin where coarse-grained glacial and fluvial sediments prevail, which exhibit high hydraulic conductivity (estimates are 0.07–0.51 cm/s in the area of the main production wells [Kleinfelder Inc., 2000]). In addition, the extraction of water from the production wells may lead to younger ages by drawing shallow groundwater deeper into the aquifer. However, results from well MW-PJOW indicate that Group 1, seasonally recharged waters already dominate the aquifer upgradient and outside the influence of the production wells. The production wells were sampled on different dates, and for three of the four production wells, mean apparent groundwater ages increased from <1 year for the April sampling to 3 years (SVPSD well 5), 4 years (SVPSD well 3), and 6 years (SV MWC well 2) when sampled during later summer months (Table 2). By contrast, ages in production wells producing older groundwater did not change between sampling dates. This finding suggests that during the dry season, the seasonal recharge (0–1 year) has moved through the aquifer and the wells draw in water that recharged in prior years, having taken longer flow paths to well capture zones.

[35] The vigorous flow system sampled by wells in Group 1 is recharged on a short time scale (≤1 year) and over a limited spatial extent (given the short time period for saturated zone transport). The production wells in this group have screened intervals from 11 to 20 m long at depths of 11–39 m but do not reach depths near the bedrock basement (Figure 7). Although the creek is not likely a major source of recharge to these wells (based on the excess air and carbon isotopes results in this study), the predominance of young groundwater in the alluvial aquifer suggests that it is this young component that likely provides much of the potential base flow to the stream. Under climate change scenarios with earlier snowmelt and runoff, this groundwater reservoir will be depleted earlier, providing less base flow and possible extreme low flows in the creek during summer and fall.

Figure 7.

Schematic cross section through Olympic Valley running along Squaw Creek, showing the approximate locations of the major faults and the depth to bedrock as based on well logs [West-Yost and Associates, 2003] and seismic profiling [Gasch and Associates, 1973]. Solid vertical lines represent screened intervals for wells on or close to the cross section, with labels indicating the mean apparent tritium-helium groundwater age (year), percentage of pre-modern water, and noble gas recharge temperature (°C). A curved line separates the wells in Group 1 with recent recharge from wells in Group 2 that are screened into bedrock and sample water with longer flow paths. Vertical exaggeration is 5.

[36] Several of the lower valley monitoring wells and the two horizontal wells (Figure 8) exhibit somewhat older 3H-3He apparent ages. The apparent groundwater ages calculated for these wells give the groundwater age histogram its bimodal character (Figure 8). These wells are grouped with wells dominated by relatively recent recharge because they do not share dissolved gas characteristics associated with bedrock groundwater. The wells do not produce groundwater containing radiogenic 4He, but they have higher concentrations of tritiogenic 3He than samples with <1 year ages (Figure 6). In contrast to the production wells with <1 year mean ages, many of these wells are screened in lower permeability media, which includes near surface fine-grained sediments in the lower valley and near surface fractured rock (in the case of the two horizontal wells). This older component may contribute to stream base flow in the lower reaches of Squaw Creek, given its occurrence in shallow monitoring wells adjacent to the creek. Significant flow and deep pools are observed in downstream reaches of the creek later in the water year than in upstream reaches.

Figure 8.

Histogram of apparent groundwater ages.

4.3.2. Long Flow Paths

[37] The remaining production wells (SVPSD Well 1, SVPSD Well 2), well T4, MW 5D, and five of the valley monitoring wells fall into Group 2, drawing a component of significantly older groundwater as evidenced by a concentration of radiogenic 4He greater than 2 × 10−9 cm3 STP/g and high pre-modern fractions (Table 2). These wells produce mixed aged water (21%–98% pre-modern), as they also all contain tritium and have mean apparent groundwater ages (for the portion of the water containing tritium) of less than 50 years. All of these wells tap the deeper flow system associated with bedrock that underlies the alluvial fill, either being partially screened in bedrock or being situated near a major fault (Figure 7).

[38] Granitic rocks have comparatively high U and Th concentrations, which can result in a relatively high radiogenic 4He production rate [Andrews et al., 1989]. In addition, glacial tills and weathered granites have been shown to exhibit high 4He release rates into circulating groundwater [Beyerle et al., 1999; Van der Hoven et al., 2005]. Nonetheless, radiogenic 4He concentrations in affected Olympic Valley wells are low in comparison to production wells affected by crustal He in bedrock wells elsewhere [Holocher et al., 2001; Manning and Caine, 2007] and in deep supply wells elsewhere in California [Hudson et al., 2002; Moran et al., 2002; Moran et al., 2005]. This component is not observed in many of the production and monitoring wells screened exclusively in alluvium, which may be a reflection of the shorter residence time for water in the alluvium. Alternatively, the crustal fluid may not be produced within the alluvium but rather may be related to diffusion of 4He from low permeability bedrock at the base of the alluvial aquifer. Although present at depth, and clearly affecting wells that directly tap bedrock groundwater or are affected by focused flow along faults (Figure 7), this component is minor in comparison to the very young groundwater component and is not likely to play a significant role in stream interaction or base flow to the stream.

4.3.3. Deep Upwelling Fluids

[39] Samples that fall into Group 3 have smaller components of recent recharge, along with an older component containing crustal He, and a component of dissolved gases from a magmatic source. Recently, Kulongoski et al. [2005] and Saar et al. [2005] have presented methods for quantifying mixing proportions for groundwaters that have crustal, magmatic, and tritiogenic components, as revealed by examination of the isotopic composition of dissolved helium. Compared to samples from those studies, Olympic Valley groundwater samples have much smaller magmatic and crustal components. The 3He in Olympic Valley samples is predominantly from atmospheric equilibrium, dissolved excess air, and from the decay of 3H. However, because the magmatic 3He/4He ratio is drastically different from the 3He/4He ratio for other helium sources, the presence of a small component of magmatic helium in four of the samples makes determination of a 3H-3He age highly unconstrained (age labeled as “3He excess” in Table 2). These samples lie along a transect that lines up with an active fault, Valley Fault 3 (Figure 1 and Figure 7), and are clearly affected by gases that emanate from a deep, magmatic source. One of these samples, MW 329, contains less than 1 pCi/L 3H and thus is a mixture of only magmatic and crustal components. A simple linear mixing calculation for this sample (using 3He/4Hemagmatic = 1.22 × 10−5 [Graham, 2002] and 3He/4Hecrustal = 6 × 10−7 (Lawrence Livermore National Laboratory, unpublished data, 2010)) results in estimates for the magmatic component of only 7%. The other three wells along Valley Fault 3 contain tritium, and tritiogenic 3He, and are complex mixtures of relatively recent recharge, older water with a significant radiogenic 4He component, and magmatic gases that reach shallow groundwater via the active fault. One additional monitoring well sample from the lower valley (MW302) is likewise affected by magmatic He.

4.4. Determination of Recharge Sources Based on DIC

[40] Water samples from the horizontal and production wells contain 16–32 mg/L C as DIC with δ13C-DIC values that range from −19.1‰ to −16.4‰. The lower valley monitoring wells have δ13C-DIC values that range from −20.9‰ to 4.2‰, with concentrations from 20 to 222 mg/L C. Stream waters are low in DIC concentration (3–13 mg/L C) and have δ13C values that range from −11.0‰ to −2.6‰.

[41] The carbon isotope values for the production well groundwater samples (δ13C-DIC of −19.1‰ to −16.4‰) are consistent with the incorporation of soil CO2 during recharge [e.g., Blumhagen and Clark, 2008; Cerling et al., 1991], which reflects a mix of respiration CO2 and atmospheric CO2 sources. The incorporation of soil CO2 in the production well groundwater suggests that recharge occurs in subalpine areas with developed soils such as the vegetated slopes surrounding the valley, as opposed to the bare rock exposures that are prevalent over the highest elevation areas surrounding the valley.

[42] When plotted against 1/[DIC], the δ13C-DIC values indicate mixing between three dominant sources (Figure 9): (1) groundwater recharged through the soil zone, (2) recharge from Squaw Creek, and (3) upwelling of deep groundwater containing carbon derived from a magmatic source. The compositions of these end-members are discussed below.

Figure 9.

Stable carbon isotope compositions (δ13C-DIC) and inverse concentration of DIC for wells, a spring, and creek water in the Olympic Valley study area. Mixing lines are plotted between two potential end-members for groundwater recharge through soils, a magmatic water end-member, and a stream end-member.

4.4.1. Recharge Through the Soil Zone

[43] The concentration and isotopic composition of soil CO2 vary in relation to the respiration rate [Cerling et al., 1991]. Consequently, recharge through soils with a range of soil respiration rates might be expected to result in a range of DIC concentrations and isotopic compositions. For this reason we selected two potential end-members for groundwaters recharged through the soil zone. In both cases, wells with moderate groundwater ages (11–17 years) were selected to avoid wells that might receive significant recharge from the creek. The average values of [DIC] (29 mg/L C) and δ13C-DIC (−17.5‰) for the Horizontal Wells were selected as the first groundwater end-member. The Horizontal wells are influenced by recharge through soils over higher elevations than the valley monitoring wells. Monitoring well MW-5D was selected as the second groundwater end-member, and has an average [DIC] of 17.2 mg/L C and an average δ13C-DIC value of −19.1‰. These end-members will be used to constrain recharge sources discussed below.

4.4.2. Magmatic Carbon

[44] The monitoring wells have a very broad range of DIC concentrations and isotopic compositions. Many of the monitoring wells have DIC compositions that are similar to those observed in the production wells. However, wells along Valley Fault #3 have much higher concentrations of DIC (50–222 mg/L C) and appear to be influenced by a carbon source with δ13C values close to −5‰. Similar δ13C values and high [DIC] have been linked to upwelling of magmatic fluids and seismic activity along faults around Mammoth Mountain in California [Sorey et al., 1998]. A contribution from a magmatic source is also consistent with the 3He/4He ratios discussed previously. Well MW-330 was chosen as an end-member to represent groundwater that has interacted with a magmatic source of fluids. Well MW-330 was selected because it has the highest [DIC] of 222 mg/L C and has no detectable tritium (<1 pCi/L).

4.4.3. Recharge From Squaw Creek

[45] The concentrations and isotopic compositions of DIC observed in the stream water samples are generally consistent with DIC derived from equilibration with atmospheric CO2, which has a δ13C value of approximately −8‰. Some of the stream δ13C-DIC values are slightly higher than the δ13C value of air, suggesting that perhaps some of the stream DIC is derived from mineral weathering. An end-member for Squaw Creek was defined by the average value of samples collected at the lower ends of the southern and northern tributaries to Squaw Creek, above the confluence (sites 1 and 2 in Figure 1), which have an average [DIC] of 4.6 mg/L C, and an average δ13C-DIC value of −4.8‰.

[46] The downstream sampling sites tend to have lower δ13C-DIC values and higher concentrations than the upper parts of the creek, which likely indicates an influx of groundwater along the stream channel (also indicated by water level and stream gauge observations, as noted above). A more detailed study quantifying the groundwater influx to the stream is ongoing and will not be discussed further here.

[47] On the basis of the end-members defined above, Squaw Creek does not appear to be a dominant source of recharge for most of the wells sampled in this study. The lone exception is the monitoring well MW-PJOW, which was the uppermost valley well sampled. On the basis of the samples collected in May and August, the [DIC] and δ13C-DIC in MW-PJOW appears to be derived from between 50% and 70% creek water. The low recharge temperatures calculated for MW-PJOW may indicate that the creek recharge occurs during the cold, high-flow conditions during snowmelt runoff. As an alternative, well MW-PJOW may receive recharge from the prominent rock escarpment to the west. Recharge through bare rock fractures would have essentially identical [DIC] and δ13C-DIC values to creek water.

[48] In general, the production wells have much higher [DIC] and much lower δ13C-DIC values than Squaw Creek. This contrast between the isotopic compositions and concentrations of DIC in the production wells and DIC in stream water suggests that these waters have different sources and indicates that Squaw Creek is not a dominant source of recharge to the production wells. However, based on the end-members defined above, samples from production wells SVPSD Well # 1, 2, and 3 fall within the range of [DIC] and δ13C-DIC values that potentially indicate between 10% and 30% of produced water is from Squaw Creek. Quantifying such small contributions of the stream water is highly uncertain due to a strong dependence on the defined compositions of end-members. Given the older apparent ages of SVPSD Wells 1 and 2, it is unlikely that they receive significant recharge from the creek. As discussed above, the slightly higher δ13C-DIC values may also indicate the presence of water that was recharged over bare rock. On the basis of the mixing analysis of DIC, the majority of water from the production wells recharged through a soil zone where respiration was active. It is unlikely that an unsaturated zone between the creek and the water table could impart the carbon isotope compositions of the production wells, since CO2 respiration would not take place. The carbon isotope results demonstrating a lack of significant recharge from Squaw Creek are in agreement with the interpretation of excess air and recharge temperature results and with stream and well hydrograph observations.

4.5. Implications of Predicted Changes in Climate

[49] Climate models predict that our study site in the Sierra Nevada Mountains is likely to see a decrease in the percentage of precipitation that falls as snow; earlier onset of snowpack melting; an increase in the number of rain on snow events; and changes in humidity, air temperature, and soil moisture [Dettinger and Cayan, 1995; Howat and Tulaczyk, 2005; Maurer and Duffy, 2005; Melack et al., 1997]. It is still uncertain whether these changes will be accompanied by a decrease in total precipitation [Hayhoe et al., 2004]. Data from our study illuminate two important aspects of the hydrological system in Olympic Valley both of which will have a direct impact on how this system responds to climate change: (1) that recharge to the alluvial aquifer occurs primarily on the lower slopes of the catchment and (2) that deep groundwater in the western part of the aquifer is very young. In this section we review the evidence for these conclusions and consider their potential impact on the response of the hydrologic system to climate change effects.

4.5.1. Recharge on Lower Slopes

[50] Groundwater sampled for this study is primarily recharged on the lower elevation slopes that surround the alluvial valley where soils are thicker and vegetation is common, rather than in upper elevation areas of exposed bedrock. Evidence for lower elevation recharge comes from NGRTs, excess air concentrations, carbon isotope compositions, and groundwater ages in the deeper alluvial aquifer. NGRTs are similar to MAATs and much higher than would be expected for direct recharge of snowmelt through fractured rock. Excess air concentrations are relatively low compared with other mountainous areas where recharge in fractured rock has been shown to lead to high excess air entrapment. Carbon isotope compositions and concentrations of DIC in the groundwater also indicate recharge through a vegetated soil zone where soil respiration is active. These carbon isotope compositions are not consistent with recharge through exposed rock or with recharge from Squaw Creek. Finally, groundwater ages in the western portion of the basin are very young; younger than ages found in horizontal wells drilled into bedrock at higher elevation. The short residence time for much of the groundwater precludes a significant contribution of recharge to these wells from distant high-elevation locations.

[51] Under the current seasonal snowmelt scenario most of the water is transported during a short snowmelt period in the spring, and much of the potential infiltration is diverted to overland flow as recharge areas become saturated. With increased temperatures, the snow line is likely to move to higher elevations in the catchment. If we assume that total precipitation remains the same, then this would mean more rain for the lower parts of the catchment where most recharge occurs. Therefore, an increase in the snow line elevation could alter the infiltration mechanisms for a significant portion of the recharge areas that feed the valley aquifer. Infiltration would change from a seasonal peak during snowmelt to a more episodic pattern spread over a longer time period during winter rainstorms.

[52] The impact of an increasing snow line elevation on recharge to the aquifer will depend to some extent on the precipitation rate during these winter storms. Winter storms with higher precipitation rates than the current snowpack melting rates are likely to cause an increase in runoff at the expense of infiltration. Whereas, precipitation rates that are lower than the current snowpack melting rate (approximately 20–50 mm/d) would spread the infiltration events out over a longer time period, providing the opportunity for more of the precipitation to infiltrate rather than run off as overland flow. From Figure 2, it appears that under current conditions, the melt rate of the snowpack generally outpaces the accumulation of precipitation as snow water equivalent during the winter. If winter precipitation were to fall as rain rather than snow, but at a similar amount and rate, it could lead to more infiltration. However, other factors such as an earlier growing season, which might increase evapotranspiration could counteract an increase in recharge. Furthermore, rain on snow events would magnify the precipitation rate and would likely exceed the current snowmelt rates over a short period of time, leading to increased overland flow and decreased infiltration.

4.5.2. Young Groundwater

[53] Groundwater ages from this study show that near the production wells, the top 10–40 m of the aquifer is dominated by seasonal snowmelt. Once the snowmelt season is over and water levels drop, these wells capture older water that infiltrated in previous years. With a decrease in infiltration, this shift would likely happen earlier resulting in an earlier onset of decreased creek flow and groundwater availability in that portion of the basin. An extremely dry winter would impact that year's water budget with little buffering by groundwater storage in the aquifer from previous years. For the same reasons, earlier onset of snowpack melting would lead to an earlier drop in water levels with little buffering from recharge of previous years. In this scenario, the detrimental effect of little to no flow in the late summer on fish and other fauna would be almost immediate. Even the moderately older groundwater ages (average MW age of 17 years) observed in the eastern portion of the basin offer only a decadal time period during which the downstream portions of Squaw Creek may be buffered against predicted changes in runoff by groundwater inflow.

[54] A decrease in total precipitation would undoubtedly lead to less groundwater recharge and less overland flow. However, because of the seasonal nature of flow in the aquifer, such changes would very quickly affect the availability of groundwater and the groundwater contribution to streamflow. This scenario is analogous to the changes observed between the very wet 2006–2007 El Niño year and the following 3 years with much lower winter precipitation (Figure 2). In 2006, the snowpack persisted for approximately 1 month longer than it did in the following 3 years. Likewise, the water level in SVPSD Well 2 remained higher than the Squaw Creek bed elevation for approximately 1 month longer in 2006 than in the following 3 years.

5. Summary and Conclusions

[55] Dissolved gas tracers provide a powerful toolset to evaluate the vulnerability of high-altitude aquifers to climate change impacts because they address key questions about recharge location and subsurface residence time. The extent to which individual catchments are vulnerable to climate change will depend largely on the specifics of geology, topography, and climate. The dissolved gas toolset used for this study can be applied under a wide range of potential settings but is especially useful in high-elevation areas because of the steep gradients in precipitation and temperature and because active recharge and vigorous flow results in relatively young apparent groundwater ages, dateable by the tritium-helium method.

[56] NGRTs constrain the location and timing of recharge and correspond to air temperatures at the time of snowmelt. Recharge occurs mainly through soil zones where the water incorporates CO2 from respiration and recharging water is thermally equilibrated within the unsaturated zone. Recharge through fractures and recharge from the creek is of lesser importance in this catchment. Predicted climate change effects such as an increase in snow line elevation and an increase in rain-on-snow events will cause the greatest impacts to the accumulation of snowpack and timing of recharge at the lower elevations of Olympic Valley. The alluvial aquifer is therefore highly susceptible to changes in climate because, as we have shown, the lower elevation areas are important for groundwater recharge. Long-term monitoring of recharge temperatures may provide a means to gauge watershed response to climate changes such as an earlier onset of snowmelt and an increase in mean air temperature.

[57] Young groundwater dominates the most permeable part of the alluvial aquifer in Olympic Valley and likely accounts for much of the potential base flow to Squaw Creek. This groundwater has an apparent age of less than one year and is therefore vulnerable to climate change over short time scales. Mixed age components also need to be considered in studies of alpine and subalpine groundwater residence time. In this study the bedrock aquifer underlying the valley fill contributes an older component that has accumulated radiogenic 4He. In addition, helium and carbon isotopes show the influence of magmatic fluids in shallow groundwater, especially in the area of an active fault.

[58] The major findings with respect to groundwater residence times in Olympic Valley, i.e., that the alluvial aquifer experiences rapid flushing of seasonal recharge and that significantly older fluids are found at the bedrock interface that underlies the alluvium are similar to the major findings of Beyerle et al. [1999] in the Lisenthal aquifer of Switzerland and Plummer et al. [2001] in the Blue Ridge Mountains. The changes in recharge, groundwater availability, and streamflow due to predicted climate change that are outlined here are likely for other small mountain catchments. Similar studies are needed in much larger mountain watersheds to determine whether an increase in scale may decrease the dominance of young waters.


[59] The authors wish to gratefully acknowledge assistance with field sampling and sample analysis by Brad Esser, Sarah Roberts, Darren Hillegonds, Mike Sharp, and Carl Gustafson. Well access and logistical support were provided by Squaw Valley Public Services District, Friends of Squaw Creek, The Resort at Squaw Creek, Squaw Valley Mutual Water Company, and Derrik Williams (HydroMetrics LLC). Funding for this work was provided by LLNL Laboratory Directed Research and Development, Climate Initiative. Jean Moran received support from the Joan Sieber research award at California State University, East Bay. We are grateful to three WRR reviewers whose comments led to improvements in the text. This work was performed under the auspices of the U.S. Department of Energy by Lawrence Livermore National Laboratory under contract DE-AC52-07NA27344.