Chemical and stable isotopic models for boundary creek warm springs, southwestern Yellowstone National Park, Wyoming

Journal of Volcanology and Geothermal Research, 43 (1990) 133-157

133

Elsevier Science Publishers B.V., Amsterdam

Chemical and stable isotopic models for Boundary Creek warm springs, southwestern Yellowstone National Park, Wyoming W.T. PARRY and J.R. BOWMAN Department of Geology and Geophysics, University of Utah, Salt Lake City, UT 84112, U.S.A. (Received August 10, 1989; revised and accepted March 1, 1990)

Abstract Parry, W.T. and Bowman, J.R., 1990. Chemical and stable isotopic models for Boundary Creek warm springs, southwestern Yellowstone National Park, Wyoming. J. VoIcanol. Geotherm. Res., 43: 133-157. Thermal springs of the Boundary Creek hydrothermal system in the southwestern part of Yellowstone Park outside the caldera boundary vary in chemical and isotopic composition, and temperature. The diversity may be accounted for by a combination of processes including boiling of a deep thermal water, mixing of the deep thermal water with cool meteoric water and/or with condensed steam or steam-heated meteoric water, and chemical reactions with surrounding rocks. Dissolved-silica, Na +, K + and Ca 2+ contents of the thermal springs could result from a thermal fluid with a temperature of 200 ± 20 oc. Chloride-enthalpy and silica-enthalpy mixing models suggest mixing of 230 °C, 220 mg/1 C1- thermal water with cool, low-C1- components. A 350 to 390°C component with C1- > 300 mg/1 is possibly present in thermal springs inside the caldera but is not required to fit observed spring chemical and isotopic compositions. Irreversible mass transfer models in which a low-temperature water reacts with volcanic glass as it percolates downward and warms, can account for observed pH and dissolved-silica, K +, Na +, Ca 2+ and Mg 2+ concentrations, but produces insufficient C1- or F - for measured concentrations in the warm springs. The ratio of aNa./aH . and C1- are best accounted for in mixing models. The water-rock interaction model fits compositions of acidsulfate waters observed at Summit Lake and of low-C1- waters involved in mixing. The cold waters collected from southwestern Yellowstone Park have 5D values ranging f r o m - 118 to - 145 per mil and 51SO values of -15.9 to -19.4 per mil. Two samples from nearby Island Park have 5D values of - 1 1 2 and - 1 1 4 per mil and 5180 values of -15.1 and -15.3 per mil. All samples of thermal water plot significantly to the right of the meteoric water line. The low C1- and variable 5D values of the thermal waters indicate isotopic compositions are derived by extensive dilution with cold meteoric water and by steam separation on ascent to the surface. Many of the hot springs with higher 6D values may contain in addition a significant amount of high-D, low-Cl-, acid-sulfate or steam-heated meteoric water. Mixing models, C1- content and isotopic compositons of thermal springs suggest that 30% or less of a deep thermal component is present. For example, the highest-temperature springs from Three Rivers, Silver Scarf and Upper Boundary Creek thermal areas contain up to 70% cool meteoric water and 30% hot water components, springs at Summit Lake and Middle Boundary Creek spring 57 are acid-sulfate or steam-heated meteoric water; springs 27 and 48 from Middle Boundary Creek and 49 from Mountain Ash contain in excess of 50% acid-sulfate water; and Three Rivers spring 46 and Phillips could result from mixing hot water with 55% cool meteoric water followed by mixing of acid-sulfate water. Extensive dilution by cool meteoric water increases the uncertainties in quantity and nature of the deep meteoric, thermal component.

0377-0273/90/$03.50

© 1990 -- Elsevier Science Publishers B.V.

134

W T PARRY AND J.R. BOWMAN

boundary. Large flows of thermal waters occur in topographic lows near surface streams such as the Firehole River in the Upper and Lower Geyser Basins, whereas fumaroles and acid springs with little water discharge occur in topographically high areas such as the Madison Plateau (Fig. 1). Truesdell and Fournier (1976) suggested t h a t the diverse compositions of surface thermal waters may be accounted for by steam loss, mixing with cool near-surface water, and chemical reactions with rocks as deep ther-

Introduction The. Yellowstone National Park area is recognized as one of the most extensive geothermal regions in the world. Pleistocene volcanism has produced extensive pyroclastic volcanics, lava flows and caldera collapse. Hot springs, geysers, pools and other thermal features described by Allen and Day (1935) and Truesdell and Fournier (1976) occur over an area of nearly 3000 km 2 within and just outside the caldera

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Fig. 1. Location map of Boundary Creek thermal areas in southwestern Yellowstone Park. The Yellowstone caldera boundary is approximately coincident with the continental divide. Abbreviations are ~L = Summit Lake; S J = Smoke Jumper; UBC = Upper Boundary Creek; M B C Middle Boundary Creek; S O = Silver Scarf; L B C = Lower Bouncary Creek; B F = Bechler Ford; PF, GF, F F = Three Rivers; M A = Mountain Ash; H = Hillside.

CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS, YELLOWSTONE N.P.

mal waters flow upwards toward the surface. A large body of chemical data and careful interpretation indicate that most thermal waters of Yellowstone are mixed waters (Truesdell and Fournier, 1976; Truesdell et al., 1977; Pearson and Truesdell, 1978). The solutions thought to be involved in mixing consist of: (1) a deep, large-volume, relatively homogeneous, meteoric hydrothermal fluid at temperatures as high as 340 to 370 °C containing 300 or more ppm C1- (Truesdell and Fournier, 1976) with a long residence time (Pearson and Truesdell, 1978); (2) a shallow, cooler meteoric water that percolates downward, reacts with the rocks as it warms, and has a residence time of 1 to 20 years; and (3) a thermal water that results from condensation of steam and other volatiles from deeper boiling fluids into shallow aquifers in which hydrogen sulfide becomes oxidized by atmospheric oxygen to produce sulfate and hydrogen ions, and other dissolved constituents that are the result of reaction with surrounding rocks. All these water components are derived ultimately from meteoric water. In surface thermal springs, the deep meteoric-hydrothermal component is often extensively diluted. However, the importance of this high-temperature component at depth is demonstrated by stable isotope studies of paleo-meteoric hydrothermal systems thought to be eroded remnants of caldera systems analogous to Yellowstone (Forrester and Taylor, 1977; Criss and Taylor, 1983; Criss et al., 1984; Larson and Taylor, 1986). These studies demonstrate that anomalous 5D and 580 values occur over very wide areas indicating that the lateral dimensions of these high-temperature meteoric-hydrothermal systems are 25 to 50 k m and depths are 5 to 7 km. High heat flow values outside the Yellowstone caldera range from 100 to 300 mW m -2 with a transition, 5 to 10 km inside the mapped boundary of the caldera, to very high values of 600 to 700 mW m -2 (Morgan et al., 1977). These heat

135

flow studies suggest that thermal areas within the caldera ring fracture system probably result from a structurally and topographically controlled convection system 200 to 800 m deep with a deeper, hotter, fluid-circulation system postulated by Morgan et al. (1977) to underlie the shallow system. Lower heat flow outside the caldera does not require the deep fluidcirculation system at temperatures of 340 to 370°C described by Truesdell and Fournier (1976) and Morgan et al. (1977) inside the caldera. Hydrothermal alteration minerals observed in drill holes in rhyolites within the caldera include chalcedony, quartz, kaolinite, illite-smectite, celadonite, mica, chlorite, albite, K-feldspar, carbonates, fluorite, pyrite, goethite, opal-CT, cristobalite, mordenite and clinoptilolite (Honda and Muffler, 1970; Keith and Muffler, 1978; Keith et a l , 1978; Bargar and Beeson, 1981, 1984; Bargar and Muffler, 1982). Although no drill hole information is available for the southwestern part of Yellowstone Park, the shallow alteration assemblages in rocks of this area are likely be similar. Moderate- to high-temperature thermal springs occur in four major groups along the Boundary Creek drainage near the western boundary of the Park. Similarity of geological environment and water compositions suggest these springs result from geochemical processes similar to thermal systems elsewhere in the Park. These areas of thermal spring activity have been described and chemical and isotopic data have been collected (Hutchinson, 1980; Thompson and Hutchinson, 1980). The major objectives of the present study are to evaluate thermal springs in southwestern Yellowstone Park outside the caldera rim, to thermodynamically model spring water composition that would result from the interaction with rock, and to establish the oxygen and hydrogen isotopic compositions of water inputs into the spring system. These geochemical and isotopic techniques are then used to investigate the details of mixing, steam loss, and

136

w a t e r - r o c k reactions, and to assess the nature and sources of inputs into the Boundary Creek thermal springs. Thermal areas discussed are shown in Figure 1 and include the Summit Lake system, Smoke Jumper, Upper Boundary Creek, Middle Boundary Creek, Silver Scarf, Lower Boundary Creek, Bechler Ford, Three Rivers and Mountains Ash. An additional area included for comparison purposes is the Hillside system in the Upper Geyser Basin. Chemical compositions of additional thermal waters used in interpretations in this report are taken from Thompson and Yadav (1979), Thompson et al. (1975), Rowe et al. (1973) and Thompson and Hutchinson (1980).

Geologic and hydrologic setting Quaternary rhyolitic volcanic rocks are the dominant lithology of the area. Eruption of the volcanic rocks was punctuated by three, major, explosive eruptive events and associated caldera collapse at 2.0, 1.3 and 0.6 m.y.; the youngest rhyotites were erupted 150,000, 100,000 and 70,000 years ago (Christiansen, 1984). Geological and geophysical evidence (Smith and Braile, 1982) is used to infer that large magma bodies in the upper crust beneath the Yellowstone Plateau provided the heat that drives the hydrothermal systems. Boundary Creek flows southwards nearly parallel to the southwestern border of Yellowstone National Park and drains a large area of the Madison Plateau. The topographic and structural rim of the Yellowstone caldera is nearly coincident with the topographic divide at the crest of the plateau (Fig. 1). The topography suggests that groundwater is recharged through the rhyolitic volcanic rocks of the plateau (Morgan et al., 1977; Hutchinson, 1980) and emerges as both cool and warm springs along the Boundary Creek, Bechler River and Ash Creek drainages outside the caldera rim in southwestern Yellowstone Park. Groundwater must also flow northwards from the crest of the Madison plateau to

W T PARRY AND J.R. BOWMAN

recharge shallow systems in the Upper Geyser Basin.

Procedures Water samples were collected following procedures of Presser and Barnes (1974). Each sample was filtered through a 0.45-#m pore diameter filter. One sample was acidified to pH < 2, one sample was diluted by a factor of 10 to prevent silica polymerization, and one filtered and one unfiltered sample were untreated for anion and isotopic analysis, respectively. The samples were chemically analyzed using methods outlined in Brown et al. (1970). Temperature and pH were measured in the field. Dissolved silica and cations were measured by atomic absorption spectrometry. Sulfate was determined by an indirect atomic absorption method in which sulfate is precipitated with barium chloride and excess barium is measured. Aluminum was determined by graphite furnace atomic absorption. Chloride and fluoride were measured with specific ion electrodes, and chloride was checked by Mohr titration. Analytical results are shown in Table 1. The collection and hydrogen and oxygen isotope analyses of the cold and thermal waters were done with standard techniques (Friedman, 1953; Epstein and Mayeda, 1953; Nehring and Truesdell, 1977). The isotopic compositions are reported relative to SMOW (Craig, 1961b).

Chemical geothermometers The fundamental assumptions that are made in attempts use geochemical indicators of' subsurface temperature include the following (Fournier et al., 1974): (1) Temperature-dependent chemical reactions among solutions, rocks and minerals occur at depth and reactant supply is not a limiting factor. (2) Chemical equilibrium is attained between water and minerals within the rock at the

137

CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS, YELLOWSTONE N,P.

temperature of the deep t h e r m a l reservoir. (3) The chemical composition of the water, derived through equilibrium processes at depth, is preserved as the water flows to the collection point at the surface. Geological circumstances commonly prevent these assumptions from being fulfilled. For example, hot water coming from a deep reservoir may mix with cooler, shallow groundwater before reaching the surface, the hot water may cool by boiling and steam separation which serves to concentrate the nonvolatile constituents in the water, or the rate of chemical reaction between water and minerals may be sufficiently rapid compared with the rate of ascent of the water so t h a t the high-temperature geochemistry is not preserved. Silica The measured silica contents of the springs in the southwestern portion of Yellowstone Park are shown in Table 1 and are plotted on Figure 2. The temperatures of last equilibration with quartz were calculated using equations from Fournier (1977) assuming t h a t the spring waters cool by separation of steam or by conduction with no steam separation and are shown in Table 1. These temperatures are computed assuming t h a t no quartz precipitates during cooling while the spring water is in transit to the surface and t h a t no mixing of shallow groundwater has occurred and thus represent m i n i m u m temperatures. Temperatures range from 130 °C to 189°C except for cold springs and vapor-dominated springs. The dissolved silica could result from additional geological processes. First, the dissolved silica could result from dissolution of glass in the rhyolite volcanic rocks from which the springs flow. Second, the dissolved silica may result from dissolution of glass or silica minerals at elevated temperature followed by mixing with cool water with or without steam separation. Third, quartz may precipitate during adiabatic or conductive cooling while the t h e r m a l water is in transit to the surface. Figure 2 has been constructed to il-

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Fig. 2. Dissolved silica (rag/l) and temperatures of southwestern Yellowstone National Park spring waters with hypothetical reaction paths and mixing lines. Quartz saturation curves are calculated from equations in Fournier and Rowe (1966). Line A-B is a simple mixing line; CDEA and C'D'E'A are the result of boiling and mixing. Open circles = Upper Boundary Creek; filled circles = Three Rivers; triangles = Mountain Ash; open stars = Lower Boundary Creek; filled stars = Middle Boundary Creek; open squares = Silver Scarf; square/star = Summit Lake; circle/star = Bechler Ford.

lustrate each of these processes. Rates of dissolution of glass and precipitation of quartz are calculated using the rate model and rate constants of Rimstidt and Barnes (1980). Curve I in Figure 2 shows the increase in dissolved silica t h a t would result from meteoric water percolating into fractures in the warm volcanic rocks and dissolving the volcanic glass. Curve II represents the temperature-composition path of the same solutions t h a t encounter fractures cont a i n i n g quartz at depth in the subsurface at a

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CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS, YELLOWSTONE N.P.

temperature of 90°C. Heating continues as the water percolates downward, b u t saturation with quartz is maintained to a temperature of 190 °C. The fluid (saturated with quartz at 190 ° C) then percolates outward along lava flow boundaries or upwards along fractures and reaches the surface at hot-spring localities. As the water reaches cooler regions, it cools by conduction and steam separation and becomes supersaturated with quartz. Quartz will then precipitate. Curve IIIa and IIIb illustrate the silica concentration in the water if cooling from 190 to 80 ° C takes place in 7 hours (IIIa) and in 73 hours (IIIb). The additional complication of mixing with cool, lowsilica water is illustrated with the dashed curve IV in Figure 2 which illustrates the consequences of mixing solutions at 80 ° C on curve IIIa with a cool water. Each of these processes is a possible source of uncertainty in estimated temperature. The silica content of the thermal springs could result from a single thermal fluid with a temperature of 190 to 200 ° C from which varying quantities of silica have precipitated during cooling. Mixing with cool water may also be responsible for many of the observed silica values.

Na-K-Ca Temperatures of last equilibration of thermal waters with Na-, K- and Ca-bearing minerals were calculated using the equation of Fournier and Truesdell (1973) and are shown in Table 1. Such estimates are generally less sensitive to the effects of dilution than are temperatures inferred from dissolved silica because the calculated temperatures are based on ratios of ions. However, the chemical reactions that produce the ion ratios are dependent on a number of factors in addition to temperature. Steady-state chemical reactions for which a continued supply of reactant hydrogen ion is available from dissolved carbon dioxide or from oxidation of hydrogen sulfide could produce solutions with alkali metal ratios determined by the dissolution of the rocks rather than by temperature dependent exchange equilibria. Potassium-rich

139

rocks would result in high potassium contents and anomalously high Na-K-Ca temperatures. Calculated Na-K-Ca temperatures of southwestern Yellowstone P a r k thermal waters shown in Table 1 range between 87 and 192°C and are generally within 20°C of the quartz saturation temperatures as suggested by Thompson and Hutchinson (1980). The 270°C temperature inferred for the Summit Lake thermal area is anomalously high, possibly because of acid dissolution of the volcanic rocks and deviations from the basic assumptions of chemical geothermometry. This water which has a low pH, high sulfate, low bicarbonate and low chloride, may originate from a vapordominated system as described by White et al. (1971). The low chloride content of the water indicates that little if any of a deep thermal component is present, and the high sulfate and low pH probably result from oxidation of hydrogen sulfide. The high potassium content is attributable to reaction of the acid water with local volcanic rock.

Mixing models Mixing of thermal and nonthermal waters is a dominant process in the thermal systems in southwestern Yellowstone Park and in other areas of Yellowstone. Mixing and mixing models have been described and evaluated by Fournier et al. (1974) and Truesdell and Fournier (1975, 1977). Chemical and physical properties of mixed waters that may be used to evaluate the characteristics and relative abundances of the end-members include enthalpy, chloride, dissolved silica, sulfate and fluoride.

Chloride-enthalpy Chloride-enthalpy mixing models assume that a deep, chloride-rich thermal water rises toward the surface and cools by steam separation and mixing with a cool, chloride-poor water. The resulting warm springs have chloride compositions and heat contents that are the result of both boiling and mixing. The model and graphical

140

W.T. PARRY AND J,R. BOWMAN

method of interpretation are described by Truesdell and Fournier (1975). Application of this model to the thermal springs in southwestern Yellowstone Park is illustrated in Figure 3. In Figure 3, variations in spring composition caused by steam separation are shown as lines connecting steam containing negligible chloride and 2,775 J/g enthalpy with the chloride and enthalpy content of hot spring waters. Lines radial to cold spring waters CS-1, CS-2 and Boundary Creek systems represent cold-water dilution of the hot-water component. In this model the least mixed water (highest chloride) is the Bechler sample and the indicated temperature of the hotwater component is dependent upon the silica content (in terms of calculated enthalpy) chosen for the cold water end-member. Figure 3 shows several choices that may be made: the composi-

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tion labeled CS-1 represents the enthalpy computed from dissolved silica for a cold spring, CS-2 represents the measured heat content (from measured temperature) for the same spring, and the third choice is the linear trend represented by the plotted water compositions. The temperature indicated for the hot-water component is 228 °C if the linear trend is chosen, 264 ° if CS-1 is chosen, and 298°C if CS-2 is chosen. However, the temperature of the deep reservoir at Yellowstone is possibly as high as 350 to 290°C with a chloride content of 300 or more mg/1 (Truesdell and Fournier, 1976). This fluid is shown on the chloride-enthalpy plot in Figure 3. The Boundary Creek systems could have resulted from cooling of this water to 310°C accompanied by steam separation followed by mixing with cool, chloride-poor surface water. Summit Lake waters (Thompson and Yadav, 1979; this study) shown on Figure 3 could represent condensed steam or steam-heated shallow water that has reacted with volcanic rock to derive a high silica content (represented on Figure 3 as high computed enthalpy) while retaining a tow chloride content. The chloride-enthalpy mixing model illustrated in Figure 3 suggests that the Silver Scarf thermal system may contain as much as 27% of a deep thermal component.

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Dissolved silica The dissolved-silica content of thermal springs may also be affected by mixing with cooler waters as described above. The silica and enthalpy of thermal springs, for example Upper Boundary Creek, could result from mixing a cold spring water at point A on Figure 2 with a 240 ° C thermal water at point B on Figure 2. Observations at Upper Boundary Creek, Summit Lake and elsewhere suggest that steam separation is also an important process. Silica contents of other springs could result from separation of steam from a hot thermal fluid with an initial temperature of 160 to 200 °C that boils and mixes with cool water shown as dashed lines in Figure 2. For example, a thermal fluid that has equilibrated with quartz at 160 ° C that boils at

] 41

CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS, YELLOWSTONE N.P.

100°C would evolve along the line CDE. Mixing with cold water (A on Fig. 2) would result in the linear trend of compositions shown,

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Fluoride and sulfate correlate with dissolved silica (enthalpy) similar in some respects to C1(Figs. 4 and 5). Sulfate shows an additional complication t h a t has been evaluated in detail by Truesdell et al. (1978). Mixing of cool groundwaters carrying about 9 ppm dissolved oxygen with the hot subsurface waters containing dissolved H2S produces an additional 13 mg/1 sulfate by oxidation of the H2S. The original deep water contained about 11 ppm sulfate, and addition of 13 ppm sulfate by oxidation of H2S results in the 24 ppm sulfate observed in some high-chloride springs at Norris. Additional sulfate can be produced by oxidation of H2S at the surface where an unlimited supply of oxygen is available; such oxidation is responsible for the

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, 20

2i4

FLUORIDE (rag/I)

Fig. 4. Fluoride-enthalpy (J/g) comparison of southwestern Yellowstone thermal waters. Summit Lake (+) and Hillside (x) data from Thompson and Yadav (1979). See Figure 2 for key to symbols.

Some of the chemical constituents in the thermal waters of southwestern Yellowstone must have been derived through chemical reaction with the enclosing volcanic rocks. The reaction between rock and water has been modeled for comparison with spring chemistry using six basic assumptions. First, the beginning water composition is assumed to be pure rain or snow melt water. Second, the source of dissolved car-

142

W T PARRY AND J.R, BOWMAN

bon (bicarbonate in the chemical analyses of Table 1) is assumed to be derived from dissolution of gaseous CO2; hydrolysis of dissolved CO 2 provides the hydrogen ion that drives chemical reaction with the rocks. An alternative source of hydrogen ion, oxidation of H2S has not been modeled. Third, the time-temperature path of the water reacting with rocks follows an arbitrarily assumed heating rate and reaction of water with volcanic rocks. The silica content of the water is calculated from the reaction rate model and rate constants of Rimstidt and Barnes (1980). The temperature and silica content of the water is summarized in Figure 2. Alternative combinations of fracture width and times could be selected to closely approximate the observed dissolved-silica contents of the warm springs. Fourth, the constituents Na, K, Ca, Mg, A1, Silica, C1 and F in the water are assumed to come exclusively from reaction with volcanic rock with the rhyolitic chemical composition shown in Table 2. Fifth, minerals precipitate when the water reaches saturation according to the equilibrium constants computed for each temperature. Sixth, the stable minerals in the model are gibbsite, kaolinite, muscovite, paragonite, KTABLE 2 Rhyolite composition used in the numerical estimates of rock-water irreversible mass transfer Constituent

Weight percent

Number of moles of ions

SiO 2 A1203 Fe203 FeO MgO CaO Na20 K20 C1 F S

76.4 12.2 0.9 0.6 0.13 0.4 3.3 4.9 0.05 0.02 (0 to 0.15) 0.04 (0 to 0.4)

1.272 0.2393 not modeled not modeled 0.0032 0.0071 0.1065 0.1040 0.00145 0.00105 not modeled

Boyd (1961), Hamilton (1963).

feldspar, albite and quartz consistent with observations in both fossil and active hydrothermal alteration systems (Giggenbach, 1984; Fournier, 1985). Metastable phases chalcedony, opal or cristobalite were not include in the model. Mordenite, clinoptilolite, interstratified clay minerals, and smectite were not included because of lack of adequate thermodynamic data. Omission of these phases is not expected to cause large errors in model calculations of water chemistry for the following reasons. In persistent, active hydrothermal systems dissolved silica is controlled by the solubility of quartz above 90°C (Fournier, 1985). Zeolites are probably unstable with respect to quartz and feldspar (Hawkins et al, 1978). Laboratory experiments and calculations show that r o c k fluid reactions involving clays, zeolites, micas or feldspars may generate solutions with similar cation ratios (Benjamin et al., 1983; Michard, 1987; Giggenbach, 1988). The course of chemical reaction is simulated with numerical calculations using the computer program PHREEQE (Parkhurst et al., 1980, Plummer et al., 1983). Equilibrium constants for mineral saturation and aqueous speciation are given in Table 3. The procedures employed in applying PHREEQE to mass transfer calculations at southwest Yellowstone are similar to procedures described by Parkhurst et al. (1980) and Plummer et al. (1983). An irreversible reaction defined by the stoichiometry of rhyolite glass (Table 2) is added to the solution as required to reach equilibrium with gibbsite, kaolinite, muscovite or K-feldspar as desired in the simulation. Quantities of mineral phases dissolved or precipitated to reach a specific molality of dissolved silica are also computed. In the model, irreversible mass transfer proceeds at each temperature until the model dissolved-silica value shown as curves I and II on Figure 2 is reached. The reaction path, in terms of the composition of fluids, minerals dissolved and precipitated, and rock dissolved, is dependent on the temperature path of the fluid. The reaction path from 20 to 80 °C has been modeled

143

CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS. YELLOWSTONE N.P.

T ire 4

~

3 Gibbsite

o

5 o

T

T

t

e

~-'~\~K-Feldspar

\r%6o

;

'5*',%

11 ', I I

Range of 8rou*ndry Creek cor~posi~ions I

5

I D +O z 4 O

O _~

Kaolinite

/ ,,x'l .'"

I

I

3

/..'"

/../Mi×ing~,~.

]"]

..'] / .,'/ /

Muscovite

.>" /

\ Kaolinite

90 ° 190°C I/-

,/

//

©

,,\

0

Y

React ion path

"2

t 2

/

'T Albite

(- Feldspal

\. !

0 -2 ~ -6

-5

~

-4

-3

-2

-I

!

2

4

5

LOG (aK+/aH+)

LOG aH4SiO4

Fig. 6. Logarithmic activity diagram showing reaction path for 20, 40, 60 (dashed lines), and 80 °C (heavy solid line). Large filled circles represent model silica values at 20, 40, 60 and 80°C.

Fig. 7. Logarithmic activity diagram for 90, 140 and 190°C and quartz saturation showing reaction path and mixing line compared with compositions of southwestern Yellowstone thermal waters.

so that the dissolved-silica content of the water is not restricted to quartz saturation. Solution compositions, masses of minerals precipitated and dissolved, and mass of rock dissolved are shown in Table 4, and the reaction path is plotted in Figure 6. The total cumulative mass of phases have been determined by the point in the reaction progress where the model silica values are reached. The results of model calculations are shown in dissolved silica-temperature (Fig. 2), log aH4SiO4 (Fig. 6), and log log (Fig. 7) diagrams. Solution parameters and moles of minerals precipitated are shown in Table 4. Rain and snow melt water containing CO 2 reacts with the volcanic rocks and reaches gibbsite saturation, gibbsite kaolinite equilibrium, followed by dissolution of all of the gibbsite that has precipitated and continued precipitation of kaolinite. The solution at 20 °C will contain the amount of dissolved silica shown in the kinetic model discribed above (9 mg/1) following dissolu-

tion of 53.9 mg of rock, and precipitation of 63.9 micromoles of kaolinite at a pH of 5.2. The solution continues to warm and dissolve additional rock. At 40°C the solution reaches 50 mg/1 dissolved silica following dissolution o f l 2 1 mg of rock and precipitation of 145 micromoles of kaolinite. Between 40 and 60°C the solution reaches saturation with K-feldspar and begins to precipitate that phase along with kaolinite. At 60°C, a dissolved-silica concentration of 155 mg/1 is reached following dissolution of 327 mg of rock, precipitation of 255 micromoles of kaolinite and precipitation of 276 micromoles of K-feldspar. At 80 degrees the solution reaches 275 mg/1 dissolved silica following dissolution of 541 mg of rock and precipitation of 368 micromoles of kaolinite and 559 micromoles of K-feldspar. At this stage in the model the solution is just saturated with albite. The solution then reaches a depth where quartz is present in the fractures in the rock and the dissolved-silica precipitates as additional quartz on these fractures as the solution warms

(aK+/aH+)--log (aK+/aH+)-

(aNa+/an+)

144

WT.

PARRY

AND J.R BOWMAN

TABLE 3 E q u i l i b r i u m constants used in t h e e q u i l i b r i u m and m a s s t r a n s f e r models LogK=A+BT+

C/T + D Log T + E / T 2

Reaction A

B

D

C

E

Reference

2A13+ + H20 + 2H4SiO 4 = A12Si205(OH) 4 ÷ 6H + - 1652.60 - 0.357007278 59009.8

633.910

(1)

3A1 :~* + 3H4SiO 3 + K + = KA13Si3010(OH) 2 ~ 10H + -2501.63 -0.543576702 89903.9

960.557

(1 )

3A13: + 3H4SiO 4 + N a * = NaA13Si3010(OH) 2 + 10H ~ -2531.98 -0.546477339 92724.9

970.918

(1)

A13+ + 3H4SiO 4 + N a + = NaA1Si30 s + 4H ~ + 4H20 - 827.174 -0.182993459 28624.8

318.762

(11

306.114

(1~

A13+ + 3H4SiO 4 + K + = KA1Si30 8 + 4H + + 4H20 - 791.220 -0.178264932 25951.2 H ~ + CO32- = HCO 3107.887 0.032528490 -5151.79 = H2CO 3 464.1966 0.093448130 Ca 2~ +CO32- = CaCO 3 1228.732 -0.299440 Ca 2÷ + H * + CO3 2 - = C a l i C O 3" 1317.0071 0.34546894 Mg 2~- + C 0 3 2 - = MgCO 3 32.172 Mg 2+ + H + + C032- = M g H C 0 3 ~ 48.6721 0.03252849

-38.9256

563713.9

(2)

2H* + C O 3 2

-

-26986.2

-0.016346

H4SiO 4 = H2SiO42- + 2H + 39.478 -0.065927

165.760

35512.75

485.818

--39916.84

-517.708

1093.486

12.7243

-2614.335

--18.0026

-

H4SiO 4 = H3SiO 4- + H* 6.368

-

2248629.

(2) (2)

563713.9

(2) (3)

563713.9

(4)

- 3405.900

(5)

- 12355.10

(5)

f r o m 80 t o 9 0 ° C a n d m u s c o v i t e s a t u r a t i o n is r e a c h e d . T h e r e a c t i o n p a t h for q u a r t z - s a t u r a t e d s o l u t i o n s f r o m 90 t o 1 9 0 ° C , s h o w n i n F i g u r e 7 , is based on the assumption of quartz saturation and saturation with kaolinite and muscovite at 9 0 ° C . T h e s o l u t i o n c o m p o s i t i o n is a s s u m e d to t r a v e r s e t h e m u s c o v i t e f i e l d a n d r e a c h Kf e l d s p a r s a t u r a t i o n a t 1 4 0 ° C . F r o m 140 to

190°C, saturation with K-feldspar and muscov i t e is a s s u m e d a n d t h e s i m u l a t i o n is t e r m i n a t e d w h e n a l b i t e s a t u r a t i o n is r e a c h e d . A t 90 °C, 5 3 0 0 micromoles of quartz have precipitated to reach q u a r t z s a t u r a t i o n a n d e q u i l i b r i u m is a c h i e v e d w i t h m u s c o v i t e . F r o m 90 t o I 4 0 ° C e q u a l increments of rock are dissolved to reach Kf e l d s p a r s a t u r a t i o n . E q u i l i b r i u m is m a i n t a i n e d

145

CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS, YELLOWSTONE N.P.

TABLE 3 (continued)

H4SiO 4 = SiO 2 + 2H20 H + + F = HF H + + 2F = HF 2 Ca 2. + F - = CaF + Mg 2÷ + OH = Mg(OH) ~ Ca 2. + OH = Ca(OH) + Mg 2÷ + F - = MgF + Na + + C 0 3 2 - = NaCO 3Na" ÷ H + + C O 3 2 - = NaHCO 3 A13T + H20 = AI(OH)2. + H + A13+ + 2H20 = Al(OH)2 + + 2H + A13. + 3H20 = Al(OH) 3 +3H + A13. + 4H20 = Al(OH)4- + 4H + A13+ + F - = A1F2+ A13. + 2F- = A1F2+ A13. + 3F- = A1F3 A13+ + 4F- = A1F4

log K (25 ° C)

5H (kcal)

Reference

-4.00 3.169 3.749 0.940 -11.794 -12.598 1.820 1.268 10.08 -4.99 -10.1 -16.0 -23.0 7.01 12.75 17.02 19.72

6.20 3.46 4.55 3.798 15.419 14.535 4.674 8.911 -3.604 11.9

(6) (7) (7) (7) (7) (7) (7) (7) (7) (7) (7) (7) (7) (8) (8) (8) (8)

44.06 20. 2.50

References: (1) Equilibrium constants determined at P and T along the boiling curve for pure water using equations, data, and the program SUPCRT of Helgeson and Kirkham (1974a, b, 1976), Helgeson et al. (1978, 1981). Regression coefficients were then computed using a multiple regression analysis. (2) Plummer and Busenberg (1982). (3) Siebert and Hostetler (1977a). (4) Siebert and Hostetler (1977b). (5) Ryzhenko (1967). (6) Helgeson (1969). (7) Parkhurst et al. (1980). (8) Hem (1968)

b e t w e e n muscovite, K-feldspar a n d q u a r t z as t h e solution c o n t i n u e s to w a r m to 190 ° C a n d h i g h e r . The solutions at this point m a y still r e a c t w i t h u n s t a b l e volcanic glass, b u t m a y be n e a r equilibrium with minor mineral phases within the rock. B e c a u s e volcanic rock would c o n t i n u e to dissolve, t h e r e a c t i o n p a t h w o u l d c o n t i n u e to follow the p a t h s s h o w n on F i g u r e 7 up to e q u i l i b r i u m w i t h muscovite, K-feldspar a n d albite. The m a s s e s of rock consumed, moles of m i n e r a l s precipitated, a n d solution p a r a m e t e r s are s h o w n in Table 4. Solution p a r a m e t e r s derived from the irreversible m a s s t r a n s f e r m o d e l i n g are c o m p a r e d w i t h a c t u a l compostions of t h e r m a l s p r i n g w a t e r , in

Table 5. The Na, K, silica a n d b i c a r b o n a t e corn t e n t s of Middle B o u n d a r y C r e e k a n d Silver S c a r f s y s t e m s are closely a p p r o x i m a t e d by model c a l c u l a t i o n s at 140 to 190°C, b u t A1 and M g conc e n t r a t i o n s are h i g h e r in the model. Na, K, A1, silica, C1 a n d b i c a r b o n a t e in the S u m m i t L a k e s y s t e m are closely m a t c h e d by model calculations at 80 to 140°C. C o n c e n t r a t i o n s of cations K + , N a +, Ca 2+ a n d M g 2+ a n d p H c a n be approxi m a t e d by a p p r o p r i a t e selection of model p a r a m e t e r s despite u n c e r t a i n t i e s i n t r o d u c e d by o m i t t i n g m e t a s t a b l e p h a s e s a n d p h a s e s for w h i c h t h e r m o d y n a m i c d a t a are lacking. U p p e r G e y s e r B a s i n a n d N o r r i s c o n t a i n more Na, silica, C1 a n d F t h a n a n y of the model calculations, a n d

-1.5

- 1.5

- 1.5

-1.5

1.5

-0.5

-1.5

-0.5

- 1.5

--0.5

20

40

60

80

90

90

140

140

190

190

K-feldspar Muscovite Albite

K-feldspar Muscovite Albite

K-feldspar Muscovite

K-feldspar Muscovite

Muscovite Kaolinite

Muscovite Kaolinite

Reach model mSiO 2 KaoliniteK.feldsparAlbite

Reach model mSiO2 KaoliniteK-feldspar

Reach model mSiO 2 Kaolinite

Reach model mSiO21 Kaolinite

Reaction path index

7.45

7.55

7.65

7.69

6.62

7.38

7.63

7.25

5.94

5.20

pH

1.33

- 1.33

-2.68

-2.68

-3.19

-3.19

-2.35

2.59

--3.08

-3.78

3.87

3.87

4.18

4.18

2.98

2.98

2.20

2.33

1.86

0.41

log mSiO 2 log K/H

5.22

5.22

5.07

4.85

3.81

4.23

4.38

3.72

1.87

0.40

7433.0

4970.0

3553.0

1467.0

1533.0

685.0

541.0

327.0

121.0

53.9

Rock (mg)

- 368.0

--255.0

145.0

-64.0

Kaolinite

3140.0

909.0

-2323.0

-634.0

-791.0

--112.0

Muscovite (micromoles

-3525.0

-3462.0

- 559.0

-276.0

K-feldspar

67439.0

-15239.0

--33748.0

-12251.0

-14061.0

-5300.0

Quartz

log Na/H Mass of rock a n d m i n e r a l s dissolved or precipitated

1Model mSiO 2 corresponds to the dissolved silica concentration estimated fl'om kinetics of dissolution of glass,

logPco 2

T (°C)

S u m m a r y of irreversible mass transfer calculations. Positive values correspond to dissolution a n d n e g a t i v e values to precipitation

TABLE 4

F~

CHEMICALAND STABLEISOTOPICMODELSFOR BOUNDARYCREEKWARMSPRINGS,YELLOWSTONEN.P.

147

TABLE 5 C o m p a r i s o n of chemical composition of solutions derived from t h e o r e t i c a l calculations of r o c k - w a t e r re action with a c t u a l spring compositions in mg/1 Model or Spring

T pH (°C)

Mineral

Ca

Mg

Na

K

A1

SiO 2

CI

HCO 3 F

0.5

0.001

155.0

0.15

30.0

0.06

II.0

0.13

0.04

232.0

0.24

20.0

0.09

0.4

13.0

0.15

0.02

275.0

0.28

31.0

0.11

2.0

0.5

17.0

1.6

0.2

40.0

0.4

57.0

0.1

Kaolinite Muscovite

4.0

1.2

38.0

9.5

0,04

40.0

0.8

201.0

0.3

140 7.69

Muscovite K-feldspar

4.0

1.1

36.0 13.0

8.4

124.0

0.8

86.0

0.3

140 7.65

Muscovite K-feldspar

8.0

2.1

66.0 24.0

6.6

124.0

1.00 213.0

0.5

190 7.55

Muscovite K-feldspar-Albite

14.0

3.9

122.0

9.3

119.0

279.0

2.6

17.0

1.0

190 7.45

Muscovite K-feldspar-Albite

19.0

5.2

161.0 12.0

96.0

272.0

4.00 290.0

1.3

0.1

175.0

83.00 246.0

16.00

Silver Scarf C e n t r a l Group (Thompson and Hutchinson, 1980); this report) 88 6.7 3.7 0.11 177.0 10.7 0.09

209.0

107.00 270.0

22.00

Hillside (Thompson a n d Yadav, 1979) 85 6.7

60 7.25

Kaolin K-spar

0.8

0.2

7.0

80 7.85

Kaolin K-feldspar-Albite

1.3

0.4

80 7.63

Kaolin K-feldspar-Albite

1.5

9O 7.38

Kaolinite Muscovite

90 6.62

Log(Pco 2) = -1.5 Log(Pco _) = -1.3 z Log(Pco ) = -1.5 2 Log(Pco ) = -0.5 2 Log(Pco _) = -1.5 '~ Log(Pco ) = -0.5 2 Log(Pco ) = -1.5 2 Log(Pco ) = --0.5

2

Middle B o u n d a r y Creek (Thompson and Hutchinson, 1980); t h i s report) 69 6.7 4.8 0.32 153.0 7.1

7.4

7.0

0.12

186.0

68.00 234.0

12.00

U p p e r Geyser B a s i n - E a r S p r i n g (Thompson et al., 1975) 95 9.0 0,6 0.01

335,0 16.5

0.35

362.0

417.00 174.0

26.00

N o r r i s - P o r k Chop (Thomspon a n d Yadav, 1979) 70 7.3 5.3

0.02

444.064.0

472.0

712.00

39.0

6.00

S u m m i t L a k e T h e r m a l A r e a (this report) 85 5.4

0.01

51.0 16.0

277.0

4.00

81.0

4.00

0.5

0.14

!40.0

0.6

148

higher t e m p e r a t u r e and salinity inputs are probably the cause. The concentrations of C1 and F - cannot be approximated in the model by stoichiometric dissolution of volcanic rocks; too little C1- and F - are present in comparison with other constituents. Inclusion of additional alumino-silicate mineral phases observed in drill holes will not resolve this discrepancy. E i t h e r the C1- and F - must be obtained by mixing a fraction of some water (the deep brine) containing these anions, or t hey must be preferentially extracted from rock during reaction progress. A f u r t h e r comparison of the irreversible mass t ra n s f er model with solution mixing is made on Figures 2 and 7. The variation of the ratio of Na + and K + to H + with mixing has been e v a l u a ted using the computer program SOLVEQ (Reed, 1982). In this computation, the q u a n t i t y of ionizable hydrogen is computed for the endmembers involved in the mixing. Mixing is t he n accomplished by conserving enthalpy, Na +, K ~ and ionizable hydrogen and a new ratio of activities of Na + and K + to H + is computed. The results of mixing Boundary Creek cold spring water with Bechler are shown on Figure 7 where the result of mixing is compared with observed activity ratios for the Boundary Creek systems. Measured warm-spring silica values may be represented in the mass transfer model, but mixing of high- and low-temperature waters appears to be a b etter explanation for most of the waters shown on Figure 2. The ratio aK+/aH+ is repesented very well in the mass t r ans f er and the mixing models, but the ratio of aNa e/a H ~ is best represented in the mixing model (Table 5). The low Na + contents of solutions derived from mass tr an s f er modeling do not model the more saline solution compositions. Mixing with a more saline reservoir fluid is a more plausible explanation.

Hydrogen and oxygen isotopes The hydrogen and oxygen isotope compositions of the cold springs and precipitation

w,'r

PARRY AND J.R

BOWMAN

samples are compiled in Table 6 and plotted on Figure 8. The cold waters from southwestern Yellowstone P a r k area have 5D values ranging from - 1 1 8 to - 1 4 5 per mil and 5180 values ranging from - 15.9 to - 19.4 per mil. The range of 5D values of these cold springs covers the entire range of D values observed for warm springs in southwestern Yellowstone Park. In contrast, the two precipitation samples from the Island P a r k Ranger Station (Table 6, Fig. 8) have 5D values of - 1 1 2 and - 1 1 4 per mil and 5180 values of - 15.1 and - 15.3 per mil, respectively. These waters are unaffected by evaporation or exchange with rock, as they plot, within analytical error, on the meteoric water line (Fig. 8). The hydrogen and oxygen isotope compositions of the w arm springs samples are illustrated in Figures 8 t hrough 10 and are compiled in

/" - rio

-

f20

× I30

i"

x

8D

~

×

I40

t?

:~

~1 phJltlp$ iii 5 r" Q 0 4 6 ~ 65 c~1~8"

,/ ~/×

4

o

0 "s"

$p

• tSO

- t'g

- f6

~4

Fig. 8. Plot of SD and 5180 values of precipitation, cool surface springs, and thermal spring waters from the southwestern Yellowstone Park region. The meteoric water line (Craig, 1961a) is included for reference. Symbols are X = precipitation and cool springs from Yellowstone; circled X = Island park; Thermal Springs are solid squares = Upper Boundary Creek; half-filled squares = Middle Boundary Creek; open squares = Lower Boundary Creek; open triangles = Silver Scarf; asterisk = Bechler Ford; solid circles = Mountain Asb; open circles = Three Rivers; cross = Summit Lake. Data from Tables 1 and 6.

CHEMICALANDSTABLEISOTOPICMODELSFORBOUNDARYCREEKWARMSPRINGS,YELLOWSTONEN.P.

149

TABLE 6 Hydrogen and oxygen isotope compositions of cold surface springs and precipitation, Southwest Yellowstone Park Sample

Location or date

5D

51SO

Big Springs Bechler

Targhee National Forest Precipitation collector 10/12/84

-138 -144

-18.5 -19.4

Snow Creek

Above confluence with Robinson Rd. #092, Targhee Na- - 1 3 3 tional Forest

-17.7

Snow Creek

Off Forest Rd # 094. Targhee National Forest

-138

-18.6

Snow Creek Fish Creek Springs

BMW ROS, Targhee National Forest -135 Targhee National Forest: 3 km south of Horsefly Springs - 1 3 7

-18.2 -18.2

Horsefly Springs

Targhee National Forest: Near intersection~f National Forest Road 082 and Fish Creek

-132

-17.8

Moss Springs

Targhee National Forest

-135

-18.3

Bear Springs

Intersection of National Forest Roads 549 and 082, Targhee National Forest

-135

-18.0

Trail Springs

On Bechler River Trail, - 4 km NE of Iris Falls

-122

-16.4

Boundary Creek

Junction with Bechler River Trail

-135

-17.4

Robinson Creek

Junction with West Boundary Trail

- 136

- 18.2

Warm River Springs

Targhee National Forest, 13 km south of intersection of national Forest Roads 112 and 150, on road 150.

-134

-18.1

Warm River

Targhee National Forest, at location above

-128

-17.4

Bechler

Precipitation collected 10/28/84

-118

-15.9

Cold Spring

Along Boundary Creek

138

-18.6

9 Ag

Campsite, Bechler River Trail

-138

-18.2

Cold Spring

Bear Junction of Shoshone Lake and Bechler River Trails near Grants Pass

-144

-19.0

Island Park Ranger Station

Precipitation collected 10/13/84

-112

-15.1

Island Park Ranger Station

Precipitation collected 10/14/84

-114

-15.3

Boundary Creek

Precipitation collected 9/14/84

-132

-17.5

Fish Creek Along Forest Rd 092

Targhee National Forest

-131

T a b l e 1. All s a m p l e s plot s i g n i f i c a n t l y to t h e r i g h t of t h e m e t e o r i c w a t e r line (Fig. 8), inc l u d i n g t h e cold s p r i n g (14°C) in t h e U p p e r Boundary Creek Thermal Area. The total range of SD ( - 145 to - 129 p e r m i d a n d t h e 5180 ( - 18.3 to - 1 5 . 4 per mil) v a l u e s is l a r g e , b u t t y p i c a l of t h e isotopic v a r i a t i o n s o b s e r v e d in o t h e r geotherm a l a r e a s in Y e l l o w s t o n e P a r k , such as t h e Nor-

ris G e y s e r , L o w e r G e y s e r a n d S h o s h o n e G e y s e r B a s i n s (Truesdell et al., 1977). H y d r o g e n a n d oxy g e n isotope c o m p s i t i o n s of w a r m s p r i n g s v a r y c o n s i d e r a b l y in m o s t of t h e i n d i v i d u a l geotherm a l a r e a s . As local cold m e t e o r i c w a t e r s in t h e w e s t - s o u t h w e s t e r n p a r t of Y e l l o w s t o n e N a t i o n a l P a r k h a v e 5D v a l u e s of - 1 4 0 to - 1 4 5 per mil (Table 6 a n d T a b l e 1, t h i s report; T r u e s d e l l et al.,

150

1977), some of these waters have experienced significant enrichment in D as well as 180. Several of the spring waters are distinctive from the main group in their isotope composition and water chemistry. Summit Lake area-A and Upper Boundary Creek-57 are characterized by high 6D and the highest 6180 values recorded, extremely low C1- contents (4 mg/1) and low pH (5.4). These could be classified as acid-sulfate waters, although the SO 4 content of Upper Boundary Creek-57 is not high. Alternatively, this sample may be steam-heated meteoric water. The water sample from Bechler Ford has a low 5D value (-145 per mil) but a high 5180 value (-15.9 per mil), and 260 mg/1 C1- suggesting a different origin from the other springs. The measured temperature and C1 content of the spring waters are not well correlated with the 8180 values when all the thermal areas are considered together (Fig. 9). Measured temperature and C1- content correlate negatively with 180 in the Three Rivers thermal area. The two samples from this area with the highest measured temperature (M and 3) have surprisingly low 5180 values, while the cooler springs have progressively higher 8180 values and generally lower C1- contents. These inverse relationships are opposite those normally expected from the processes of surface evaporation (Craig et al., 1963) or progressive interaction between water and enclosing rock (Craig et al., 1956; Craig, 1963; Clayton et al., 1968). Too few samples were analyzed in each of the remaining thermal areas to establish such correlations. The 8D and C1- values for the springs are plotted on Figure 10. These parameters are not well correlated for all the thermal areas together, but increasing AD values are approximately correlated with decreasing C1- contents in the Three Rivers thermal area. This relationship is also inconsistent with that expected from surface evaporation effects.

Discussion o f the isotope results The hydrogen and oxygen isotopic composi-

W,T. PARRY AND J.R. BOWMAN

tions suggest that many of the hot springs contain variable but significant amounts of three separate components: (1) a cold, low-chloride meteoric water; (2) an inferred hot, high-chloride meteoric water; and (3) a deuterium-rich, chloride-poor component that is likely an acid sulfate-steam condensate or steam-heated meteoric water. The acid-sulfate water and cool meteoric water components are both of local origin and comprise the major fraction of water

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CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS, YELLOWSTONE N.P.

in the observed hot springs. The ultimate specific recharge area for the hot meteoricthermal component is not indicated by the isotope data. Truesdell et al. (1977)have demonstrated that the processes of mixing and boiling can be traced using C1 and D because C1- is a conservative component in geothermal systems below 300°C and because changes in the isotopic composition of hydrogen are sensitive to the mechanism of boiling (continuous vs single-stage steam separation). We have used their approach in our analysis of the springs from southwestern Yellowstone as shown in Figure 10 which is analogous to Figure 8 of Truesdell et ah (1977). The temperature and C1- contents of the deep water (point I, Fig. 10) selected for this analysis are those indicated by the chloride-enthalpy characteristics of the spring waters, 298 ° C, 175 rag/1 (Fig. 3). Other end-members are possible if the actual enthalpy of the cold water endmember is ignored. Selection of point I (Figure 10) provides a m a x i m u m contribution from the thermal water component in any mixture of the cold and hot meteoric-water components (Curve A, Fig. 10). Using an alternate, highertemperature and C1 -thermal-water endmember requires even greater dilution by cold meteoric water and the high-SD, low-C1- water component (acid-sulfate or steam-heated water) to explain the measured 5D and C1- values of the hot springs (see later discussion). The 5D value of the hot meteoric water end-member, - 145, is the value measured for one of the very few cold springs existing in the area (Upper Boundary Creek cold spring) and is consistent with 5D values of - 1 4 5 per mil for several of the hot springs. For comparison, cold meteoric waters in the Norris Geyser Basin have 5D values of - 1 4 0 to - 1 4 4 (Truesdell et al., 1977). The effects of boiling this thermal end-member to 93°C were computed using the equations of Truesdell et al. (1977) and are shown as curves B and C on Figure 10. Only the spring at Bechler Ford plots close to either to these curves (C), and is consistent with a process of continuous steam

151

separation. Thus it is the only one of the spring waters sampled that could have evolved directly from the nearly undiluted thermal water component. The very much lower C1 values of the other hot springs require extensive dilution of the thermal water with low-C1 water. The effects of mixing the thermal water (point I) with cold meteoric water (T = 14°C, 5D = - 1 4 5 , C1 = lmg/1) are shown as curve A (Fig. 10). The final 5D and C1- contents of these mixtures after boiling to 93°C are shown by curves D and E for single-stage and continuous steam separation, respectively. The temperatures in parentheses along these curves are those of the water mixture prior to boiling. Only the compositions of the two highest-temperature waters from the Three Rivers thermal area (3 and M), the two hightemperature springs from the Silver Scarf thermal area (65 and 138), and the spring S from the Upper Boundary Creek thermal area are within or close to the range of 5D and CI-- values defined by this two-component, mixing/boiling model. The D and C1 contents of these five samples indicate dilution of hot water by 50 to 70% by mass with cold meteoric water. At this degree of dilution, it is not possible to distinguish the mechanism of boiling. Most of the spring samples, including the high-temperature springs from Middle Boundary Creek thermal area (27 and 48), plot at significantly higher 5D values and lower C1 contents than the 5DCl-field defined by boiling of the meteoricwater/deep-thermal water mixtures. These discrepancies suggest mixing of a third water component, enriched in D and greatly depleted in CI-, with the boiled meteoric-water/deepthermal-water mixtures. One possibility is relatively D-enriched cold meteoric water such as those of Boundary Creek (-132) and associated springs. Another reasonable candidate for this third component would be descending acid-sulfate water, such as sample 57 from Middle Boundary Creek or preferably A in the Summit Lake thermal area. Acid-sulfate waters with very low C1- contents and 5D values

152

W T PARRY AND J.R. BOWMAN

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in Figure 8. Also shown are: Curve A representing D and C1- compositions and temperatures of subsurface mixtures of cold meteoric water (T = 14°C, C1- = 1 mg/1, &D = -145) and deep thermal water (T = 298°C, C1- = 175 mg/1, designated as point I); Curve B representing compositions of water in equilibrium with steam (single-stage steam separation) boiling from 298 to 93°C; Curve C representing compositions of water boiling from 298 °C to 93 °C with continuous steam separation; Curve D representing compositions of the mixed waters of curve A after boiling in equilibrium with steam (singlestage steam separation) to 93 ° C; Curve E representing compositions of the mixed waters of curve A after boiling to 93 °C with continuous steam separation; Curve Frepresenting compositions and temperatures of mixtures of deep thermal water with the acid-sulfate water (sample SLTA-A) of Summit Lake Thermal Area (T = 85 °C, C1- = 4 mg/1, 6D = -129, designated as point A); Curve G representing compositions of mixtures of cold meteoric and acid-sulfate water; Curve H representing the compositions of the mixed waters of curve F after boiling in equilibrium with steam (single-stage steam separation) to 93°C; Curve I representing composition of the mixed waters of curve F after boiling with continuous steam separation to 93°C. Temperatures of the water mixtures prior to boiling are shown in parentheses along curves D, E, H and I. Curves A - I were calculated utilizing the equations formulated by Truesdell et al. (1977).

h e a v i e r t h a n - 1 3 0 p e r m i l also exist in N o r r i s G e y s e r B a s i n (Truesdell et al., 1977). M i x t u r e s of acid-sulfate w a t e r (point A) w i t h deep t h e r m a l w a t e r a n d cold m e t e o r i c w a t e r a r e i n d i c a t e d by c u r v e s F a n d G in F i g u r e 10, respec-

tively. I n t h e t h r e e - c o m p o n e n t model, s p r i n g s 27 a n d 48 f r o m Middle B o u n d a r y C r e e k t h e r m a l a r e a a n d s p r i n g 49 f r o m t h e M o u n t a i n Ash therm a l a r e would c o n t a i n m o r e t h a n 50 m a s s percent acid-sulfate w a t e r . Most of t h e s p r i n g

CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS, YELLOWSTONE N.P.

waters can be explained by a combined process of boiling of thermal meteoric water, variably but extensively diluted by cold meteoric water (50% dilution), followed by mixing with variable amounts of acid-sulfate water. The higher 5D values of the springs in the Middle Boundary Creek and Mountain Ash thermal areas are consistent with more extensive mixing with surfacederived, D-enriched acid-sulfate waters. All but two samples (Middle Boundary Creek 45 and 48) require at least some component of cold meteoric water. Many of the samples (those with 5D values less than - 1 3 9 per rail and C1 contents less than 80 mg/1) are dominated by cold meteoric water. Spring 45 from the Middle Boundary Creek thermal area plots outside the 5D-C1 field defined by these three water components, suggesting that it may represent a spring that evolved from boiling to 93 ° C of a binary mixture of 35% deep thermal water and 65% acid-sulfate water. Three Rivers springs 46 and Phillips plot outside the three-component mixing triangle in Figure 10. These are the coolest, have the lowest C1- contents, and have significantly higher 6180 contents than the two hottest springs sampled in the area. These trends are all consistent with three-component mixing, but with an acid-sulfate component with a lower 5D value than point A. The chemical-isotopic trends suggest that Three Rivers springs 46 and Phillips could result from boiling to 93 ° C of a binary mixture of meteoric and deep water (55% meteoric), followed by mixing of an acid-sulfate water with a 5D values of about - 1 3 8 per mil (similar to acid-sulfate sample Middle Boundary Creek-57. The amount of acid-sulfate water inferred as a component of the hot springs in Southwestern Yellowstone Park would be reduced if the cold meteoric water component has a 5D value higher than - 1 4 5 per rail. There are significant variations in 5D values of cold meteoric water within other parts of Yellowstone Park (Truesdell et al., 1977), but the scarcity of cold springs in the southwestern portion prevented us from documenting significant variation in this part of

153

the park. If cold meteroric waters with a 5D value as heavy as - 1 3 8 were mixed with the inferred thermal water end-member in the Middle Boundary Creek thermal area, then the hot springs 27 and 48 could be explained by single-stage steam separation at 93°C of a binary of cold and hot meteoric water mixture (60-70% cold meteoric water). Little or no contribution of acid-sulfate water would be required. Further definition of variations in hydrogen isotopic composition of cold meteoric water in Yellowstone Park is required to demonstrate that the higher 5D values of the springs in the Middle Boundary Creek and Mountain Ash thermal areas are the result of more extensive mixing with surface-derived acid-sulfate waters. Conclusions Geothermometry, mixing models, mass transfer models and isotope data show that the thermal springs in southwestern Yellowstone Park are composed of mixed waters. The endmembers involved in the mixing are: (1) a deep, hot, high-chloride, thermal meteoric component, (2) a shallow, cooler component that is locally derived; and (3) condensed steam or steam-heated water also locally derived. Silica and Na-K-Ca thermometers suggest a m a x i m u m temperature of 200 ± 20°C; mixing models indicate T _ 298°C and 220 mg/1 C1for the deep component. A low-temperature water reacting with volcanic glass as it percolates downward and warms, may achieve the concentrations of silica, K +, Na +, Ca 2+, Mg 2+ and H +, but not C1- or F - observed in the springs; a mixing component of high-C1 water is required. The specific recharge areas of the deep meteoric component are unknown, but the major volumetric contribution (70% or more) to the thermal springs is local meteoric water that has percolated into the volcanic rocks, become heated, and reacted with the rocks. This meteoric water did not come from some distant

154

geographic location such as Island Park or the Gallatin Mountains. The chloride and isotopic composition of the thermal springs suggests that 30% or less of a deep thermal meteoric component is present. Our geochemical evidence suggests that a deep thermal component with temperature as high as 360°C and chloride content of 300 mg/1 or more as has been suggested for the major geyser basins within Yellowstone Park is possible but not required to account for chemical and isotopic characteristics of the Boundary Creek warm springs outside the caldera. All of the observed thermal spring characteristics may be accounted for if the high temperature component has a temperature as low as 180 degrees and no higher than 300°C. The data can be accounted for by the presence of a smaller component of higher temperature water but such a component is arbitrary and not required in the models proposed. The low C1- contents and variable 5D values of the hot spring waters indicate that they likely have been extensively diluted with cold meteoric water and subjected to steam separation on their ascent to the surface. Further, many of the hot springs have 5D values too high and C1 -~ contents too low to be derived solely by binary water mixing and steam separation unless the cold meteoric water end member has 5D values of - 1 3 8 per mil or higher. The chemical and isotopic discrepancies suggest alternatively that many of the hot springs contain significant but variable amounts of a third, high 5D-low C1 water component, likely acid-sulfate or steamheated meteoric water. Assignment of highertemperature thermal water end-members than the ones chosen to illustate the model does not change these conclusions but does increase the relative amounts of the two diluting cold-water end-members. The lack of clear geochemical evidence for a meteoric-hydrothermal water component at T > 350°C in the hot springs of southwestern Yellowstone contrasts with the geochemical evidence for such a high-temperature compo-

W.¢. PARRY AND ,JR. BOWMAN

nent in the hot springs of the Norris and Upper Geyser Basins studied by Truesdell and Fournier (1976) and Truesdell et al. (1977). Such a high-temperature component is likely to exist beneath the Yellowstone system, at least within the thermal core, by analogy with isotopic evidence from exhumed paleo-hydrothermal systems. Identical geochemical and isotopic techniques were utilized in both these hot springs areas to infer mixing end-members. The different hot-water end-members inferred for these two areas may indicate differences in geological controls on these different hydrothermal systems in Yellowstone. The hot spring systems at the Upper Geyser Basin studied by Truesdell et al. (1977) and Truesdell and Fournier (1976) are within the caldera; hot spring systems in southwest Yellowstone studied here lie outside the caldera. Because of this difference in geological environment, it should be no surprise that the thermal water end-member outside the caldera could be derived from a relatively cooler, shallower and smaller hydrothermal system than the system from which the thermal water end-member inside the caldera was derived. The dramatic difference in heat flow inside and outside the caldera boundaries (Morgan et al., 1977) supports such a possibility.

Acknowledgments Financial support was provided by National Park Service Contract PX-1570-4-1794 and the University of Wyoming-National Park Service Research Center Grant No CX-12001B028. F r a n k DeCourten, Steve Boss, and David Susong assisted with collection of water samples. Ray Lambert performed the hydrogen isotope analyses. E.U. Petersen, P.N. Wilson, D. Hederley-Smith, R.E. Criss and one anonymous journal reviewer reviewed the manuscript.

References Allen E.T. and Day, A.L., 1935. Hot springs of the Yellowstone National Park. Carnegie Inst.

CHEMICAL AND STABLE ISOTOPIC MODELS FOR BOUNDARY CREEK WARM SPRINGS, YELLOWSTONE N,P.

Washington Publ. 466, 525 pp. Bargar, K.E. and Beeson, M.H., 1981. Hydrothermal alteration in research drill hole Y-2, Lower Geyser Basin, Yellowstone National Park, Wyoming. Am. Mineral., 66: 473-490. Bargar, K.E. and Beeson, M.H., 1984. Hydrothermal alteration in research drill hole Y-6, upper Firehole river, Yellowstone National Park, Wyoming. U.S. Geol. Surv., Prof. Pap. 1054-B, 24 pp. Bargar, K.E. and Muffler, L.J.P., 1982. Hydrothermal alteration in research drill hole Y-11 from a vapordominated geothermal system at mud Volcano, Yellowstone National Park, Wyoming. Wyo. Geol. Assoc. Guideb., Thirty-Third Annu. Field Conf., pp. 139-151. Benjamin, T., Charles, R. and Vadale, R., 1983. Thermodynamic parameters and experimental data for the Na-K-Ca geothermometer. J. Volcanol. Geotherm. Res., 15: 167-186. Boyd, F.R., 1961. Welded tufts and flows in the rhyolite plateau of Yellowstone Park, Wyoming. Bull. Geol. Soc. Am., 72: 387-426. Brown, E., Skougstad, M.W. and Fishman, M.H., 1970. Methods for collection and analysis of water samples for dissolved minerals and gases. U.S. Geol. Surv., Techniques of Water-resources Investigations, Book 5, Chapter AI: 160 pp. Christiansen, R.L., 1984. Yellowstone magmatic evolution: Its bearing on understanding largevolume explosive volcanism. In: Explosive Volcanism: Inception, Evolution, and Hazards. Nat. Acad. Sci., Washington D.C., pp. 84-95. Clayton, R.N., Muffler, L.J.P. and White, D.E., 1968. Oxygen isotope study of the River Ranch No. 1. Well, Salton Sea geothermal field, California. Am. J. Sci., 266: 968-979. Craig, H., 1961a. Isotopic variations in meteoric waters. Science, 133. 1702-1703. Craig, H., 1961b. Standard for reporting concentrations of deuterium and oxygen-18 in natural waters. Science, 133: 1833-1834. Craig, H., 1963. The isotopic geochemistry of water and carbon in geothermal areas. In: E., Tongiorgi, (Editor), Nuclear Geology on Thermal Areas, Spoleto. Consiglio Nazionale della Ricerche, Laboratorie de Geologia Nucleare Pisa, pp. 17 -53. Craig, H., Beato, G. and White, D.e., 1856. Isotopic geochemistry of thermal waters. Proc. 2nd Conf. Nuclear Processes in Geologic Settings, National Research Council, Nuclear Sci. Ser. Rept., 19: 29-38. Craig, H., Gordon, L. and Horibe, Y., 1963. Isotopic exchange effects in the evaporation of water. J. Geophys. Res., 68: 5079-5087. Criss, R.E. and Taylor, H., P., Jr., 1983. An lSOfl60

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and D/H study of Tertiary hydrothermal systems in the southern half of the Idaho Batholith Criss, Geol. Soc. Am. Bull. 94: 640-663. Criss, R.E., Ekren, E.B. and Hardyman, R.F., 1984. Casto ring zone: a 4,500 km 2 fossil hydrothermal system in the Challis volcanic field, Central Idaho. Geology, 12: 331-334. Epstein, S. and Mayeda, T.K., 1953. Variations of oxygen-18 content of waters from natural sources. Geochim. Cosmochim. Acta, 4: 213-224. Forrester, R.W. and Taylor, H.P., Jr., 1977.18Ofl60, D/H and 13Cfl2C studies of the Tertiary igneous complex of Skye, Scotland. Am. J. Sci., 227: 136-177. Fournier, R.O., 1977. Chemical geothermometers and mixing models for geothermal systems. Geothermics, 5: 41-50. Fournier, R.O., 1985. The behavior of silica in hydrothermal solutions. In: B.R. Berger and P.M. Bethke (Editors), Geology and Geochemistry of Epithermal Systems, Rev. Econ. Geol., Econ. Geol., The Economic Geology Publishing Company, E1 Paso, TX, pp. 45-61. Fournier, R.O. and Rowe, J.J., 1966. Estimation of underground temperatures from the silica content of water from hot springs and wet-steam wells. Am. J. Sci., 264: 685-697. Fournier, R.O. and Truesdell, A.H., 1973. An erapirical Na-K-Ca geothermometer for natural waters. Geochim. Cosmochim. Acta, 37: 1255-1275. Fournier, R.O., White, D.E. and Truesdell, A.H., 1974. Geochemical indicators of subsurface temperature-Part 1. Basic assumptions. J. Res. U.S. Geol. Surv., 2: 259-262. Friedman, I., 1953. Deuterium content of natural water and other substances. Geochim. Cosmochim. Acta, 4: 89-103. Giggenbach, W.F., 1984. Mass transfer in hydrothermal alteration systems - a conceptual approach. Geochim. Cosmochim. Acta, 48: 2693-2711. Giggenbach, W.F., 1988. Geothermal solute equilibria. Derivation of Na-K-Mg-Ca geoindicators. Geochim. Cosmochim. Acta, 52" 2749-2765. Hamilton, W., 1963. Petrology of rhyolite and basalt, northwestern Yellowstone Plateau. U.S. Geol. Surv., 475-C: C78-C81. Hawkins, D.B., Sheppard, R.A. and Gude, A. J., Jr., 1978. Hych'othermal synthesis of clinoptilolite and comments on the assemblage phillipsiteclinoptilolite-mordenite. In: L.B. Sand, and F.A. Mumpton, (Editors), Natural Zeolites. Occurrence, Properties, Use. Pergamon Press, New York, NY, pp. 337-343.

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Helgeson, H.C., 1969. Thermodynamics of hydrotherreal systems at elevated temperatures and pressures. Am. J. Sci., 267: 729-804. Helgeson, H.C. and Kirkham, D.H., 1974a. Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures. I. Summary of the thermodynamic/electrostatic properties of the solvent: Am. J. Sci., 274: 1089-1198. Helgeson, H.C. and Kirkham, D.H., 1974b. Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures. II. Debye-Huckel parameters for activity coefficients and relative partial molal properties: Am. J. Sci., 274: 1199-1261. Helgeson, H.C. and Kirkham, D.H., 1976. Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures. III. Equation of state for aqueous species at infinite dilution. Am. J. Sci., 276: 97-240. Helgeson, H.C., Delany, J.M., Nesbitt, H.W., and Bird, D.K., 1978. Summary and critique of the thermodynamic properties of rock-forming minerals. Am. J. Sci., 278A: 229 pp. Helgeson, H.C., Kirkham, D.H. and Flowers, G.C., 1981. Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures. IV. Calculation of activity coefi ficients, osmotic coefficients, and apparent molal and standard and relative partial molal properties to 5 kb and 600°C. Am. J. Sci., 281: 1249-1516. Hem, J.D., 1968. Graphical methods for studies of aqueous aluminum hydroxide, fluoride and sulfate complexes. U.S. Geol. Surv., Water Supply Pap. 1827B, 33 pp. Honda, S. and Muffler, L.J.P., 1970. Hydrothermal alteration in core from research drill hole Y-l, Upper Geyser Basin, Yellowstone National Park, Wyoming. Am. Mineral., 55: 1714-1737. Hutchinson, R.A., 1980. Boundary creek thermal areas of Yellowstone National Park I: Thermal activity and geologic setting. Geotherm. Resour. Council Trans., 4: 129-132. Keith, T.E.C. and Muffler, L.J.P., 1978. Minerals produced during cooling and hydrothermal alteration of ash flow tuff from Yellowstone drill hole Y-5. J. Volcanol. Geotherm. Res., 3: 373-402. Keith, T.E.C., White, D.E. and Beeson, M.H., 1978. Hydrothermal alteration and selfsealing in Y-7 and Y-8 drill holes in northern part of Upper Geyser Basin, Yellowstone National Park, Wyoming. U.S. Geol. Surv., Prof. Pap., 1054-A: A1 -A26. Larsen, P.B. and Taylor, H.P., Jr., 1986. An oxygenisotope study of w a t e r - r o c k interactions in the

W.T. PARRY

AND

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BOWMAN

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