geochemistry and dynamics of the yellowstone national park...
TRANSCRIPT
Ann. Rev. Earth Planet. Sci. 1989. 17:13-53
GEOCHEMISTRY AND
DYNAMICS OF THE
YELLOWSTONE NATIONAL PARK HYDROTHERMAL SYSTEM!
Robert O. Fournier
US Geological Survey, Menlo Park, California 94025
INTRODUCTION
The magmatic-hydrothermal system at Yellowstone National Park is unique among presently active systems in regard to its unparalleled geysers, tectonic environment, and magnitude of caldera-forming volcanic activity. However, over geologic time, systems like Yellowstone have been common and appear to have been responsible for the formation of many important basc-metal and precious-mctal ore deposits. For these reasons, Yellowstone has received a great amount of scientific attention. Geologic investigations in Yellowstone began with the Hague expedition in the late 1 800s, and the results of the first chcmical analyses of thermal waters from the Park's many hot springs and geysers were published in 1 888 by Gooch & Whitfield ( 1 888). Since then, numerous investigators have studied the hydrothermal system and its geologic environmcnt, and data collectcd over a span of a hundred years are available for comparison and interpretation.
THE GEOLOGIC AND GEOPHYSICAL SETTING
The geologic setting and geophysical characteristics of the Yellowstone National Park region have been described by many investigators. Particular attention is called to summaries by Smith & Christiansen ( 1 980),
1 The US Government has the right to retain a nonexclusive royalty-free license in and to any copyright covering this paper.
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14 FOURNIER
Christiansen (1 984), Smith & Braile ( 1984), Fournier & Pitt ( 1985), and the Yellowstone National Park Task Group ( 1 987). Yellowstone Park is situated on a high volcanic plateau that sits astride a region of active major crustal extension. This plateau and the adjacent region to the west have been the most seismically active area of the Rocky Mountains in historic time (Smith & Braile 1 984). According to Christiansen ( 1984), volcanism in the Yellowstone region has been active for the last 2 .2 Myr, and the erupted lavas and tuffs have been predominantly rhyolitic and subordinately basaltic. Major caldera-forming eruptions of silicic pyroclastics occurred about 2.0, 1 . 3, and 0.6 Ma (Figure 1 ), and each of these was preceded and succeeded by large eruptions of rhyolitic lava. The last eruption of rhyolitic lava was about 70,000 yr B .P. (Christiansen 1 984).
A variety of geophysical anomalies coincide with the 0.6-Ma Yellowstone caldera that, taken together, indicate that the magmatic system at Yellowstone is still vigorous and that temperatures in excess of 350°C exist at very shallow depths. These anomalies include low densities (Blank & Gettings 1974), large convective and conductive heat flows (Fournier et al 1 976, Morgan et al 1977, White 1 978), a magnetic low and shallow calculated Curie-point isotherms (Smith et al 1 974, Bhattacharyya & Leu 1975, Eaton et al 1 975), low seismic velocities (Iyer et al 198 1 , Daniel & Boore 1 982);high seismic attenuation (Benz & Smith 1 984, Brokaw 1 985), a sharp increase in electrical conductivity at a depth of about 5 km, shown by magnetotelluric soundings (Stanley et al 1 977), and a lack of seismic focal depths deeper than about 3-4 km beneath much of the Yellowstone caldera (Fournier & Pitt 1 985). In addition, there has been a spatial and temporal pattern of historical deformation-inflation followed by deflation (Pelton & Smith 1 979, 1 982, Jackson et a11 984, Hamilton 1 985, Dzurisin et al 1 986, Meyer & Lock 1 986, Dzurisin & Yamashita 1 987)that is difficult to explain by processes that do not involve magmatic activity.
The combined geologic and geophysical data support a model in which magma or partially molten material underlies much of the O.6-Ma Yellowstone caldera at depths ranging from about 5 to 1 0 km, and possibly as shallow as 3 km beneath portions of the eastern half of the caldera (Eaton et a1 1975, Iyer et a1 1 98 1 , Lehman et a1 1 982, Smith & Braile 1 984, Benz & Smith 1 984, Brokaw 1 985) . A sharply bounded sei,smic low-velocity zone in the western part of the Park was reported by Lehman et al ( 1 982) that would have been compatible with the presence of partially molten rock at a depth of 3-4 km in a N-S belt just to the west of Upper Geyser Basin and extending to Shoshone Geyser Basin. However, Lehman et al ( 1982) cautioned that this low-velocity zone might be fictive, the result of
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YNP HYDROTHERMAL SYSTEM 1 5
\ '-\.,"'- ,
( '., ,-�. -. '/ '-- MH����----------" / . 1 ''-. \ • . 'TB ) I f.q,. \'-" .�... "-
\ a: )CS ',\
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\ '\. . \
\. ./ . ..:) ,--. ,f'
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10 20 30 40 50km
Figure 1 Locations of major rivers draining Yellowstone National Park (dot-dashed lines) and calderas of the Yellowstone Plateau volcanic field: northeast-slanting lines within the 2.0-Ma first-cycle caldera, northwest-slanting lines within the 1 .3-Ma second-cycle caldera, and horizontal lines within the O.6-Ma third-cycle caldera. Crosshatched areas show regions of caldera overlap. Letter codes show river names and approximate locations of places mentioned in text: BC, Boundary Creek; CH, Crater Hills; CS, Calcite Springs; FAR, Falls River; FHR, Firehole River; GAR, Gardner River; GBR, Gibbon River; GC, Grand Canyon; HL, Heart Lake; HLB, Heart Lake Basin; HSB, Hot Springs Basin; JC, Josephs Coat Hot Springs; LB, Lower Basin; LL, Lewis Lake; LR, Lamar River; MB, Midway Basin; MH, M ammoth Hot Springs; MJ, Madison Junction; MR, Madison River; MV, Mud Volcano; NB, Norris Basin; RS, Rainbow Springs; SB, Shoshone Basin; SJ, Smoke Jumper; SL, Shoshone Lake; SR, Snake River; TB, Tower Bridge; UB, Upper Basin; WS, Washburn Hot Springs; WT, West Thumb; YL, Yellowstone Lake; YR, Yellowstone River. Outline of Yellowstone National Park shown for reference. Figure revised and redrawn after Christiansen (1984).
an attenuation problem in which very weak first arrivals were not detected. Subsequent seismic experiments with better recording-station coverage have shown that, indeed, there was an attenuation problem with the earlier data, and that a 4.0 km s -I low-velocity zone does not exist at a shallow level beneath the western part of the Park (Brokaw 1 985).
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1 6 FOURNIER
DISTRIBUTION AND TYPES OF HYDROTHERMAL ACTIVITY
The distribution of active and recently active hydrothermal features in Yellowstone is shown in Figure 2. They are mostly within the O.6-Ma Yellowstone caldera, close to the outer edge of the main ring fracture zone or at the margins of two resurgent domes. Tn addition, pronounced hydrothermal activity occurs in a linear band along a system of faults extending northward from Norris Geyser Basin through Mammoth Hot Springs.
o 10 20 3.0 40 50 km �'----�'----'�--�.�---'�--�'
Figure 2 Map of Yellowstone National Park, Wyoming, USA, showing O.6-Ma Yellowstone caldera, two resurgent domes, active and recently active hydrothermal features (solid black), major faults (solid lines outside caldera region), and locations of selected named thermal areas (key to place names given in Figure 1 legend). Figure redrawn after Christiansen (1984).
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YNP HYDROTHERMAL SYSTEM 1 7
Representative analyses of thermal waters from major hydrothermal localities in Yellowstone are shown in Table 1 . Waters from many of the Yellowstone hot springs and geysers were collected and analyzed in the late 1 800s (Gooch & Whitfield 1 888), and a later exhaustive study was carried out in 1 925-30 by Allen & Day ( 1935). Investigators who have collected and analyzed the thermal waters and gases since 1959 include Noguchi & Nix ( 1963), Scott ( 1964), Mazor & Wasserburg ( 1965), Gunter & Musgrave ( 1 966), Schoen & Rye ( 1970), Muffler et al ( 1971), Gunter ( 1973), Mazor & Fournier ( 1973), Rowe et al ( 1973), Thompson et al ( 1975), Truesdell et al ( 1977), Craig et al ( 1978), Kim & Craig ( 1 979), Thompson & Yadav ( 1979), Kennedy et al ( 1982, 1 985, 1 987), Mazor & Thompson ( 1 982), and Truesdell & Thompson ( 1 982). Although changes in the behavior of individual hot springs and geysers are common within the Park, there has been no significant change in the range of water compositions exhibited by the hot springs and geysers for at least a century, and probably much longer.
Most of the hydrothermal activity in the western half of the Park is typical of hot-water systems in which fluid pressures steadily increase downward and the maximum temperature attainable at a given depth is given by a boiling-point curve based on hydrostatic conditions. All of the major geyser basins in Yellowstone Park, except Norris and Heart Lake, lie within the portion of the 0.6-Ma Yellowstone caldera that overlaps the 2.0-Ma caldera (Figure 1, horizontal line pattern overlaying northeast line pattern). Rhyolite flows precede the ash-flow eruptions that formed the 2.0-Ma caldera, and it is likely that other rhyolite flows partly filled this caldera prior to the ash-flow eruptions that formed the 0.6-Ma caldera. Therefore, within the portion of the 0.6-Ma caldera that overlaps the 2.0-Ma caldera, the general sequence of volcanic rocks is likely to be precaldera rhyolite flows, relatively impermeable 2.0-Ma ash-flow tuff, permeable rhyolite flows, relatively impermeable 0.6-Ma ash-flow tuff, and permeable rhyolite flows. Thus, relatively permeable rhyolite flows are present at various depths that can serve as fluid reservoirs within relatively impermeable ash-flow tuff.
In the eastern part of the O.6-Ma caldera, where it does not overlap the 2.0-Ma caldera, most of the caldera fill is likely to be relatively impermeable 2.0-Ma and 0.6-Ma ash-flow tuffs with little interlayered permeable rhyolite flow material, although one major rhyolite flow caps some of the ash-flow tuff. In this part of the caldera, vapor-dominated hydrothermal activity is prevalent in which fluid pressures remain nearly constant through a significant depth interval (White et al 1971) . This is also the part of the Park in which magma or partly molten material may be closest to the surface (Lehman et al 1 982).
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Table 1 Chemical analyses of selected thermal waters from Yellowstone National Park "r1 0
Number 3 5 10 e ::<:I
Locality Mammoth West Shoshone Heart Lower Upper Upper Grand Norris Norris § ::<:I
Hot Springs Thumb Basin Lake Basin Basin Basin Basin Canyon Basin Basin Name Y-IO Lakeshore Washtub Unnamed Ojo Punch Ear Unnamed Porcelain Cinder
drill hole Geyser Spring Caliente Bowl Spring Sevenmile Terrace Pool Hole
Sample No. Y-IO 17484 T7214 T7348 17560 17836 17956 17615 17528 17908 Date 09/13/69 10/09/74 09/72 09/73 09/29/76 Temp. eC) 70 90 81 93.5 95 94 94 91 93.5 92 pH 7.48 7.76 9.00 9.48 1-74 8.33 8.49 8.61 3.57 SiO, (mg kg-I) 88 328 366 230 312 371 534 654 329 Al 0.01 0.14 0.06 Fe 2.4 0.05 Ca 450 2.0 0.4 0.9 0_75 0.67 0.82 0.5 2.12 6.3 Mg 80 0.51 0.05 0.01 0_01 0.02 0.01 0.Q2 0.03 0.12 Na 161 408 365 382 330 420 319 366 404 346 K 69 20 16 21 10 17 27 58 81 81 Li 1.8 3.24 1.0 6.6 3.6 3.8 4.9 3.45 5.8 3.9 NH, 1.0 0.1 HeO, 997 531 406 306 246 590 146 223 47 SO, 800 55 48 100 26 19 19 85 31 147 CI 171 261 328 365 326 289 415 427 669 569 F 4.2 14.5 25.5 36 30 28 24 12.8 5.8 6 B 3.8 3.3 2.8 2.4 3-8 3.8 3.6 16 9.9
Reference' 2 3 2 4 2 2
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Table 1 (Continued)
Number 11 [2 13 14 15 16 17 18 19 20
Locality Norris Crater Upper Hot Springs Josephs Coat Josephs Coat Upper Boundary Washburn Washburn Basin Hills Basin Basin Hot Springs Hot Springs Basin Creek Hot Springs Hot Springs
Name Echinus Crater Hills Iron Unnamed Unnamed Unnamed Hillside Unnamed Unnamed Unnamed Geyser Geyser Spring group
Sample No. 17846 17804 YF448 YF451 YF452 T9-10 17930 17304 YF429 Date 09/28/78 1920s 06/22/69 06/26/69 06/26/69 09/22/73 06/22/69 Temp. Cc) 93 88 9[ 89 86 94 8 3 92 91 86 pH 3.2 3.18 3.7 2.66 1.82 9.38 8.62 7.94 8.00 SiO, (mg kg-I) 278 676 365 3[7 333 238 170 21() 247 AI 0.60 0.13 0.2 Fe
Ca 4.9 5.64 37.1 4.1 3.4 7.9 4.1 2 17.2 Mg 0.52 1.04 Trace 3.75 0.39 0.01 0.27 0.10 4.10 9.3
i Na 166 640 77 187 109 150 18[ 9.7 27.1 K 50 133 28 16 17.1 19 7.7 1 1 .2 6.5 13.7
'1i
Li 0.7 6.9 0.04 0.01 0.[8 0.83 1.2 0.1 0.1 :t: ><
NH, 4 171 57.3 22.4 0.1 0.51 270 658 0 HC03 () () 0 0 0 335 251 259 107 8.2 :;:J
0 SO, 337 566 23[ 1530 1830 24 [4.2 10 900 1950 >-l CI 108 890 0.1 0.1 5.4 72 107 51 F 5.6 27.5 1.2 0.8 3.7 12 19.1 0.1 0.5
:;c �
B 3.0 20 2.2 5.05 0.07 0.36 1.17 1.0 6.6 7.84 :> t"' Referencea 4- 4 6
VJ >< VJ >-l
• References: rI1 (I) UnpUbliShed data, USGS (R. Barnes, analyst) (4) Unpublished data, USGS (J. M. Thompson, analyst) � (2) Thompson & Yadav (1979) (5) Allen & Day (1935) (3) Thompson et al (1975) (6) Thompson & Hutchinson (1980) ......
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20 FOURNIER
Hot- Water Systems
The typical surface expression of a high-temperature hot-water system is boiling hot springs, possibly with geyser activity, that discharge significant quantities of neutral to slightly alkaline water. The famous geyser basins of Yellowstone National Park are probably the best examples in the world of this type of hydrothermal activity. These include Upper, Midway, Lower, West Thumb, and Shoshone Geyser Basins within the caldera and Heart Lake and Norris Geyser Basins just outside the caldera (Figure 2). Thermal waters tend to emerge where faults cut across the topographically low basins, and hydrothermal activity is most pronounced at the intersections of faults. The waters are mostly chloride rich (Table I, numbers 2-9) and deposit siliceous sinter that over the years has accumulated in thick mounds and terraces. However, acid-sulfate boiling pools with little or no discharge of liquid water (Table 1 , number 1 3) and mud pots occur in some of the geyser basins where upflowing boiling water is only a few meters below the land surface. In these instances, the residual liquid-water fraction generally flows laterally a relatively short distance to emerge as a boiling spring in a topographically low area. The steam that separates from boiling, alkaline-chloride waters generally contains H2S that oxidizes to H2S04 when it comes into contact with air in perched pools of groundwater. Examples of steam-heated acid-sulfate features above shallow, boiling, alkaline-chloride water are Iron Spring in the Upper Geyser Basin (Table 1, number 1 3) , the Fountain Paint Pot in the Lower Geyser Basin, and the Thumb Paintpots at West Thumb.
Hydrothermal fluids that issue from all the geyser basins cited above flow to the surface from reservoirs that have temperatures ranging from about 1 800 to 270°C (Fournier et al 1976) at minimum depths of 1 00 to 550 m (estimated using theoretical boiling-point versus depth curves for hydrostatic conditions). These reservoirs are situated within thick sequences of rhyolitic lava flows and ash-flow tuff of relatively uniform chemical composition.
Thermal waters at Mammoth are very different from those in the geyser basins. The Mammoth thermal waters flow to the surface through a thick sequence of sedimentary rocks that includes limestone, dolomite, and gypsum-hearing shales. Consequently, they are rich in bicarbonate and sulfate (Table 1 , number 1 ). Relatively high concentrations of calcium and magnesium and low silica concentrations reflect a low temperature of water-rock equilibration at Mammoth: An isothermal temperature of 73°C was measured in a 1 1 2 .8-m well, and 88°C was estimated by chemical geothermometry (White et al 1 975). At higher temperatures magnesium and calcium are removed from solution, mostly as silicates and carbonate,
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YNP HYDROTHERMAL SYSTEM 2 1
and the fluid becomes richer in sodium and potassium. I n general, thick travertine deposits, such as those found at Mammoth, are found only where limestone is in the section traversed by the thermal fluid and reservoir temperatures are less than about 100--1 20°C.
Vapor-Dominated Systems
In vapor-dominated systems, steam and other gases (mainly CO2 and some H2S) fill open fractures throughout a significant vertical extent beneath a cap that acts as a throttle, and liquid water is present mainly in pore spaces. The vapor-dominated zone essentially acts as a heat pipe, with rising steam carrying thermal energy upward and volumetrically small amounts of liquid water, formed by condensation of steam, flowing back downward. The cap at the top of the vapor-dominated zone is generally considered to be relatively impermeable rock that may support a zone of perched water that is a mixture of local meteoric water and condensed steam. Temperatures and pressures commonly change very little within the vapor-dominated zone, which (in consequence) is underpressured with respect to a normal hydrostatic system. The existence of a hydrostatically underpressured system requires either that the vapor-dominated region be perched above the surrounding lowlands in topographically high terrain or that a region of relatively low permeability surrounds the vapor-dominated zone, effectively preventing the free movement of cold groundwater into the system.
Typical surface expressions of underlying vapor-dominated systems are fumaroles, acid-sulfate boiling pools with little or no discharge of liquid water (Table 1 , numbers 14 and 1 5), mud pots, and absence of nearby alkaline-chloride thermal waters issuing at the surface. Throughout much of the relatively high plateau region in the east-central part of the Park, violently boiling, steam-heated pools of acid-sulfate water containing little chloride and with little or no visible discharge of liquid water dot barren landscapes composed predominantly of residual quartz, clays, sulfates, and native sulfur. Vapor-dominated conditions are thought to underlie much of this region (White et aI1 97 1), and a vapor-dominated system has been confirmed by research drilling at Mud Volcano (White ct al 1 975, Zohdy et al 1973). Mud Volcano is the most accessible of these localities, but other major acid-sulfate thermal manifestations are at Hot Springs Basin, Josephs Coat Hot Springs, and Washburn Hot Springs (Figure 1 ).
Within the heat-pipe zone of large vapor-dominated systems, such as The Geysers in California and Larderello in Italy, the temperature is generally about 240°C with a corresponding vapor pressure of water about 33 bars (White et al 1 97 1). In my experience, temperatures at the tops of small vapor-dominated zones that have been explored by drilling in various
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22 FOURNIER
parts of the world generally range from about 1 50° to 200°C, and thus the steam that heats the overlying fumaroles, mud pots, and acid-sulfate pools generally separates from the liquid-water fraction at a pressure of at least 5 to 1 5 bars. Therefore, upflowing steam that is rich in CO2 with minor H2S may condense underground at a temperature above 150°C and react with the surrounding rock. This reaction produces a solution rich in sodium bicarbonate, poor in chloride (only a small amount of chloride is leached from the local rock), and poor in sulfate (oxygen is not available to oxidize H2S much below the water table) (Table 1 , number 16). The spring from which sample number 1 6 in Table 1 was collected is a very small, continuously discharging, superheated spring in a region of boiling, steam-heated pools that exhibit little or no discharge of liquid water. Chemical geothermometry (Fournier 198 1 ) applied to this water indicates a temperature of water-rock reaction of about 1 80-190°C. The composition of a nearby nondischarging, steam-heated pool is shown by sample number 1 5 in Table 1 . The latter water contains abundant sulfate derived from the oxidation of H2S that is streaming through the pool, a relatively large amount of the volatile constituent ammonia, very little chloride, and ratios of cations indicative of decomposition of rock with little water-mineral equilibration in the process.
All the chloride-poor thermal waters in the eastern part of the Park contain relatively high concentrations of ammonia (e.g. numbers 14-16, Table 1) . The highest reported ammonia concentrations have been found in feebly discharging hot springs in the Washburn Hot Springs region, about 1 km north of the northern rim of the Grand Canyon of ·the Yellowstone, where Allen & Day ( 1935) found 893 mg kg- 1 NH4 and 2444 mg kg-1 S04 in one thermal water. Subsequent collections and analyses of thermal waters from this area by the US Geological Survey (USGS) confirm that the predominant dissolved constituent in the fluids is ammonium sulfate, and silica is the next most abundant dissolved constituent (Thompson et al 1 975). The amounts of dissolved ammonium sulfate in different springs are highly variable, however. Examples of thermal waters from the Washburn Hot Springs area are given in Table 1 (numbers 19 and 20). Allen & Day (1935) noted that the gas collected from this locality contains appreciable methane ( 13 .20%) and ethane ( 1 . 1 5%) and concluded that the organic gas and ammonia are probably derived from distillation of buried sediments. In general, this conclusion appears to be justified, especially in view of the occurrence of hydrocarbons at three localities in the eastern part of the Park: Tower Bridge, Calcite Springs, and Rainbow Springs (Figure 1 ). Love & Good (1 970) studied the hydrocarbons in thermal waters from these localities and from other localities in an arcuate southeastward- to eastward-trending area about
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YNP HYDROTHERMAL SYSTEM 23
1 15 km long in northwestern Wyoming, and they concluded that the source of the hydrocarbons could be Paleozoic or Mesozoic sediments that underlie the surficial volcanics.
Evidently, hot water flushes petroleum from sedimentary rocks deep underground, and mixtures of water and petroleum rise together along fractures to the base of a vapor-dominated system situated beneath Washburn Hot Springs. There, distillation at high temperature and high pressure results in steam that is rich in organic gases and for which NH3 > H2S. Where this steam condenses at shallow levels, the accompanying H2S is oxidized to sulfuric acid, which is immediately converted to ammonium sulfate by excess NH3 (A. H. Truesdell, oral communication, 1 988). This results in a water that is slightly alkaline, containing mostly dissolved ammonium sulfate (Table 1 , numbers 1 9 and 20).
The Mirror Plateau (that includes the eastern resurgent dome and Hot Springs Basin, Figure 2) and Mud Volcano regions may not be the only places within Yellowstone Park where vapor-dominated hydrothermal systems occur. Boiling, acid-sulfate pools and mud pots occur throughout an extensive area of acid-altered rock in the Smoke Jumper Hot Springs region (Figures 1 , 2), which lies astride the Continental Divide. The outer edge of the main fracture of the 0.6-Ma Yellowstone caldera is thought to pass just to the west of this area of hydrothermal activity. It is possible that relatively dilute thermal waters with high proportions of bicarbonate to chloride found in the Boundary Creek region near the southwest corner of the Park (Figure 2 and Table 1 , number 1 8), and similar waters flowing from the Hillside group of hot springs in the Upper Geyser Basin (Table 1 , number 1 7), are in part steam condensate that flowed laterally away from the topographically high Smoke Jumper Hot Springs system (Thompson & Hutchinson 1 980).
Transition from Hot- Water- to Vapor-Dominated at End of Glaciation
The surface expression of the hydrothermal system beneath the Mirror Plateau has not always been characterized by acid-sulfate activity. Old, decaying geyser mounds and soil-covered siliceous sinter deposits are common throughout this region. Such sinter deposits form only where alkaline-chloride thermal waters flow over the ground. Therefore, the hydrothermal system has changed from hot-water- to vapor-dominated since the end of the last glaciation, about 1 4,000 yr B.P.
During glaciation, fluid pressures within the hydrothermal system increased and decreased in proportion to thc thickncss of the ice cover (discussed in detail later). The increased hydrostatic head was probably sufficient to inhibit the development of a vapor-dominated system or even
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24 FOURNIER
to temporarily condense a preexisting vapor-dominated condition. After the ice disappeared, it is likely that many years were required to generate, or regenerate, vapor-dominated conditions at depth. Also during the waning stage of glacial activity, ice probably remained within the gorge of the Grand Canyon of the Yellowstone after the ice sheet had melted from the adjacent plateau. Ice filling the Grand Canyon would have kept the regional water table high, allowing alkaline-chloride hot springs to flow at relatively high elevations. As the ice in the gorge diminished in volume, the water table slowly lowered through rock that had previously been heated to very high temperatures by upflowing thermal waters. This could cause the hydrothermal activity in topographically high places to change from outflow of alkaline-chloride water to discharge of high-pressure steam, throttled by poor permeability within highly altered and silicified rocks and by perched NaHCOrrich groundwater that serves to keep a pressure lid on the system. Alkaline-chloride waters still emerge at topographically low places in the eastern part of the Yellowstone caldera, such as deep within the Grand Canyon of the Yellowstone, where rapid erosion has cut a steep gorge more than 350 m deep, and at Rainbow Springs, which is about 1.8 km NW of Hot Springs Basin and only about 50 m lower in elevation. This suggests that some of the vapor-dominated hydrothermal activity in the eastern part of the Park may be relatively shallow rooted. However, in a region of generally impermeable rock underlain by a large heat source, it is possible that there may be rapid lateral transitions from shallow, alkaline-chloride hot-water to deep vapor-dominated conditions. Gravity and P-wave velocity data are compatible with the presence either of partial melt or a very deep and extensive vapor-dominated system extending from Hot Springs Basin to beneath the eastern resurgent dome (Lehman et al 1 982, Smith & Braile 1 984).
FACTORS AFFECTING THERMAL FLUID COMPOSITIONS
Fluid-Rock Interactions
The total dissolved solids in hydrothermal fluids are controlled in great part by the availability and solubilities of salts that can be leached from the wall rocks through which the fluids flow, augmented by anions contributed from the influx of acid gases (mainly CO2 at Yellowstone). There is the possibility, however, that some salinity may be contributed by a very small amount of deep brine of magmatic origin that mixes with a large amount of dilute water of meteoric origin (Fournier & Pitt 1 985). The maximum amount of magmatic water mixed with water of meteoric origin cannot be more than about 0.2-0.4% (discussed later).
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YNP HYDROTHERMAL SYSTEM 25
The effect of rock type on fluid compositions at Yellowstone was mentioned previously with regard to the relatively low-temperature bicarbonate- and sulfate-rich waters at Mammoth that issue from sedimentary rock, compared with the chloride-rich waters that issue from rhyolitic volcanic rocks. However, at high temperatures the effect of rock type is generally greatly diminished because magnesium becomes incorporated in clays and micas and the solubilities of CaC03 and CaS04 decrease as temperature is increased, allowing waters to become relatively rich in alkali chlorides. The compositions of the alkaline-chloride waters that issue in hot springs reflect water-rock equilibration in rhyolitic rocks at estimated temperatures ranging from about 1 60° to 350°C (Fournier et al 1976,
Truesdell & Fournier 1 976, McKenzie & Truesdell 1 977). However, lead isotope data for hot-spring deposits (travertine and siliceous sinter) indicate that all the thermal waters in the Park have a component that has been in contact with sedimentary rock (Leeman et aI1977).
Cation ratios are controlled by temperature-dependent hydrolysis and base-exchange reactions involving dissolved constituents and minerals in the wall rock. Hydrothermal alteration of rhyolitic material encountered in several relatively shallow drill holes has been described in a series of reports (Honda & Muffler 1 970, Keith & Muffler 1 978, Keith et al 1978,
Bargar & Beeson 1981, Bargar & Muffler 1982). In general, glass-rich rhyolitic material has altered to a variety of zeolites, clays, micas, and hydrothermal K-feldspar. Minerals commonly found in veins include amorphous silica, chalcedony, quartz, calcite, fluorite, pyrite, iron oxides, and various clays and micas.
Decompressionai Boiling and Underground Mixing of Waters
In addition to rock-water interactions, decompressional boiling and underground mixing of different waters are important processes that affect hydrothermal fluid compositions in the hot-water systems at Yellowstone. For example, variations in chloride and bicarbonate in Upper Basin thermal waters, shown in Figure 3, result from varying degrees of conductive cooling and decompressional boiling and from mixing of different thermal waters that come from reservoirs that have only slightly different temperatures. Waters that have cooled conductively without boiling as they flowed slowly upward to the surface fall along line AC. At the other extreme, waters that have undergone maximum decompressional boiling without significant conductive cooling fall along line BD. One end-member water has an initial composition prior to boiling shown by point A (the Geyser Hill-type water of Fournier et aI 1 976), and the other has an initial composition shown by point C (the Black Sand-type water of Fournier et
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26 FOURNIER
700
600
ill -0300 ·C o ::c 200 o
100
100 200 300 400 500 600 700 Bicarbonate, mg/kg
Figure 3 Chloride vs. bicarbonate for waters from Upper Geyser Basin. Bicarbonate concentrations are adjusted for the decrease in bicarbonate and the formation of carbonate that occurs at shallow levels by the reaction 2HCOj == COj2+H20+C02• The cluster of dots labeled M represents waters that are mixtures of thermal and dilute, cold groundwaters.
al 1976). Waters that plot below line AC have been diluted by shallow groundwaters. The Hillside group of springs (triangles) may have a component of steam condensate that would account for relatively high concentrations of bicarbonate relative to chloride, as discussed previously in the section on vapor-dominated systems.
The range in chloride concentrations shown by boiling and nonboiling hot-spring waters that are geographically close together and that have similar ratios of dissolved constituents (e.g. CI/HC03) can be used to estimate the temperature of the reservoir supplying water to the springs (Fournier 1 979). This is conveniently accomplished by making use of enthalpy-chloride relations, as shown in Figure 4. All the available data for thermal waters from Geyser Hill in Upper Basin that have Cl/HC03 > 3 (in equivalents) have been plotted in this figure. Enthalpies were obtained from steam tables (Keenan et al 1 969) and are those of thc residual liquid at the measured collection temperature. In Figure 4 the initial enthalpy and chloride concentration of the reservoir fluid (point R) are given by the intersection of line SB, drawn from steam to the highest measured chloride concentraton (point B, assumed to have undergone maximum possible boiling), with the line drawn parallel to the enthalpy axis through the lowest chloride concentration (point A, assumed to have cooled entirely by conduction because of a very low rate of discharge and a temperature
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600
500
(1l .::£. "-.400 (1l E
�300 Q) -0
·c 0200
-.C U
100
0
0
0
B
YNP HYDROTHERMAL SYSTEM 27
___ �II:R
500 1000 1500 2000 Enthalpy, J/ 9
s 2500 3000
Figure 4 Enthalpy vs. chloride for waters from Upper Geyser Basin.
less than the boiling temperature). Thus, prior to decompressional boiling and/or conductive cooling, the initial reservoir fluid contained about 360 mg kg -l chloride at an enthalpy of about 920 J g - l (2 I S°C). To the extent that the naturally discharging spring waters do not cover the entire range of "maximum possible boiling" to "entirely conductively cooled" conditions, the estimated initial enthalpy and chloride in the reservoir water will be in error. Also, in the above process care must be taken to
exclude waters that have low chloride concentration because of mixing with dilute, usually cool groundwaters, such as the waters represented by the cluster of dots labeled M in Figure 3. Commonly, dilution with shallow groundwaters can be detected by increased concentrations of Mg in the resulting mixed water and by flow rates that seem too large to be compatible with conductive cooling.
Silica can be used even more effectively than chloride to estimate the temperature of the reservoir that is feeding a group of hot springs and geysers because of the constraint that the solubility of quartz controls the concentration of silica in the reservoir fluid. Figure S is an enthalpy-silica
diagram that shows the solubility of quartz (Fournier & Potter 1982) and measured silica in the Geyser Hill waters that were plotted in Figure 4. As with the enthalpy-chloride relations shown in Figure 4, the range in silica concentration plotted in Figure S can be used to determine the enthalpy
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28 FOURNIER
600
500
�400 B
"" 01
E 300 0 A
-0 u
:= 200 (f)
100 E ., Ql en
0 0 500 1000 1500 2000 2500 3000
Enthalpy. Jig Figure 5 Enthalpy vs. silica for waters from Upper Geyser Basin.
(converted to temperature with steam tables) and silica concentration in the reservoir fluid prior to boiling (point R). Point R coincides with the solubility of quartz curve and also has the same enthalpy as point R in Figure 4. The enthalpy-silica and enthalpy-chloride data present a remarkably consistent picture.
Variations in the amount of boiling and different amounts of mixing of two or more waters can be exhibited by samples collected at different times from a single spring, as shown by frcqucnt sampling and analyses of waters from Cistern Spring in Norris Basin (Fournier et al 1 986). Waters from this spring were collected and analyzed five times from May 1966 through June 1 975. During that period the chloride concentration ranged from 455 to 475 mg kg-I, except for the last sample (290 mg kg- I ) that was collected one month before a local magnitude 6. 1 earthquake. Because the 38% decrease in chloride might have been a prccursor of the earthquake, a monthly collection and analysis was made of Cistern Spring from 1 976 to 1 985 to check for possible changes in composition preceding subsequent earthquakes. The results, in terms of CI and S04, are shown in Figure 6, along with data for Echinus Geyser, Steamboat Geyser, and springs with low sulfate collected at Porcelain Terrace (all in Norris Basin). Since 1 975, no correlation with seismic activity has been found, but large changes in the chloride/sulfate ratio and the degree of boiling have been observed
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Ol .::c
400
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CD 200 E ai � �
100
00
•
� . • \ I. to. � t I \I-�
¥.� A \ c:. \
\ \---
---------
,;---
200
• Cistern Spring
• Steamboat Geyser
• Echinus Geyser, pre-1979
to. , ' 1979-1980
• Acid-sulfate seeps
" Springs with lowest sulfate
1<Oown-hole sample, Y-12
+ Calculated composition before boiling
M
. . .5
��' n·· � :\:;.:;:. • *-.!I" " ..... ,,� . -
400 600
Chloride, mg/kg
Figure 6 Chloride vs, sulfate for Cistern Spring and other selected thermal waters from Norris Geyser Basin (from Fournier et aI1986),
� 't:I
�
i > t"" '"
� tv �
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30 FOURNIER
that correlate with widespread seasonal boiling phenomena that occur in Norris Basin every August or September. In Figure 6 thc lcast boiled waters lie closest to the chloride-sulfate mixing line passing through point M' and the star indicating the composition of water collected from the Y-1 2 research well drilled at Porcelain Terrace by the USGS (White et al 1 975). More extensively boiled waters lie farther above this line. The data are interpreted as indicating mixing of three types of water: a chloride-rich and sulfate-poor component with a temperature> 260°C, a chloride-poor and sulfate-rich component with a temperature < 200°C, and a dilute, relatively cool component with little chloride or sulfate (Fournier et al 1986). Thc most pronounccd mixing (dccrease in chloride and increase in sulfate) occurs immediately after episodes of intense boiling. Variations in intensity of boiling apparently are related to changes in pressure within the hydrothermal system that are affected by several factors, including seasonal changes in water table, self-sealing by mineral deposition, clogging of channels by clay suspended in flowing waters, geyser eruptions, deformation and fracturing induced by tectonic forces, and hydrofracturing.
DISSOLVED GASES RELATIVE TO AN ENTHALPY-CHLORIDE MODEL
Enthalpy-Chloride Relations
Using chemical geothermometers, ranges in composition resulting from different proportions of boiling and conductive cooling (as in Figures 3-6), and limited information from relatively shallow drill holes, Fournier et al ( 1976) estimated chloride concentrations, temperatures, and enthalpies of the shallowest reservoirs supplying thermal waters to the various geyser basins east of the Continental Divide and plotted the data on an enthalpychloride diagram, similar to Figure 7. They concluded that all the investigated waters could be related to a deep water containing about 400 mg kg-' chloride at an enthalpy of 1 570 J g-' (about 335-340°C) by boiling (lines radial to the enthalpy of steam) and by mixing with a cold dilutewater component (lines approximately radial to the origin). Truesdell & Fournier (1 976) performed a similar analysis, including water from Shoshone Geyser Basin west of the Continental Divide, and concluded that a deep "parent" water might have a temperature as high as 350-370°C. A major underlying assumption in the construction and interpretation of Figurc 7 is that both boiling (with separation and removal of steam from contact with the residual liquid) and mixing of fluids occur deep underground as well as near the surface, and that there may be complete or partial water-rock equilibration along the flow path or within a reservoir at an intermediate depth after boiling or mixing of fluids has occurred.
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YNP HYDROTHERMAL SYSTEM 3 1
1000
900
800
0'1 � 700 CH ........... 0'1 600 E NBO
500 GB Sy Q)
"'0 400 ' ;::
0 ..c 300 U
200 MH 0
100 OHS
c.� ST 0
0 500 1000 1500 2000 2500 3000
Enthalpy, Jig Figure 7 Enthalpy of liquid water vs. chloride for various aquifers. Letter codes as follows: 7M, Sevenmile Hole; BS, Black Sand-type water; CH, Crater Hills; CW, cold water; DUB, deep Upper Basin 270°C water; GB, Gibbon; GH, Geyser Hill-type water; HL, Heart Lake Basin; HS, Hillside Springs; LB, Lower Basin high chloride/bicarbonate water; MB, Midway Basin; MH, Mammoth Hot Springs; MJ, Madison Junction; NB, Norris Basin; P, the deep parent water of Fournier et al (1976); R, deepest and hottest reservoir fluid; SS, Shoshone Basin high-chloride water; ST, steam; SY, Sylvan Springs.
The model that is presented in Figure 7, though very speculative, provides a reasonable framework for interpreting gas and isotope data.
Fournier et al (1976) suggested that 270De water might be widespread beneath Upper and Lower Geyser Basins in a permeable zone at an intermediate depth between deeper, higher temperature water and shallower 20Q-2l 5De local reservoirs that feed the hot springs and geysers. In the construction of Figure 7 this 2700e water (point DUB) was assumed to be a mixture of the deeper � 3400e water (point P) and cold, dilute water (point eW), and it was further assumed that no conductive cooling occurred before or after mixing. Therefore, the 270De water would be expected to contain all or most of the gases originally present in the deeper water (Fournier 1981 ), plus an atmospheric component of gas dissolved in the diluting water and a small amount of additional gas that might be leached from the local rock, such as radiogenic argon (Kennedy et al 1 987). In contrast, waters that have undergone considerable boiling and loss of steam moving from the deeper reservoir to a shallower reservoir, such as
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32 FOURNIER
Norris, are expected to be somewhat depleted in deep-component gases. The distribution and variations in the compositions of the major gases (primarily CO2) and the noble gases (including their isotopes) in the various Yellowstone waters are consistent with this model.
Carbon Dioxide and Bicarbonate Waters in shallow reservoirs feeding springs in Lower Basin, Midway Basin, the western and northern parts of Upper Basin, the Sevenmile Hole region of the Grand Canyon of the Yellowstone, and Shoshone Geyser Basin all cluster close to the mixing line drawn from the 270°C water (point DUB, Figure 7) to dilute water. Departures from that line can be the result of a small amount of conductive cooling and/or a small amount of boiling of waters moving from the deeper to the more shallow reservoirs. Again, the thermal waters in these shallower reservoirs would be expected to contain all or most of the gaseous components originally present in the deeper reservoirs, but modified by the gases in the diluting water and by reactions with the wall rock. In these waters, little CO2 has been lost by partitioning into a separating steam phase. Therefore, CO2 is available to form carbonic acid and react with the wall rock as temperatures decrease, resulting in higher bicarbonate concentrations (e.g. Black Sand-type water) than are present in waters that have lost CO2 by boiling at high temperature before they equilibrated chemically with rock in a lower temperature reservoir (e.g. Norris and Geyser Hill-type waters) (Fournier 1 98 1). The chloride-rich, low-sulfate, pH-neutral thermal waters at Norris have Cl/HC03 ratios > 1 3, the Geyser Hill-type waters (Figure 3, type A) have ratios ranging from about 3 to 8, and the Black Sand-type waters have ratios ranging from about 0.8 to 1 .0.
Noble Gases and Their Isotopes The early studies of the relative abundances of dissolved gases, including Ne, Ar, Kr, and Xe, in the thermal waters of Yellowstone (Mazor & Wasserburg 1 965, Gunter & Musgrave 1 966, Mazor & Fournier 1 973, Gunter 1 973) were interpreted as indicating that these gases originate largely from infiltrating meteoric water saturated with atmospheric gases. However, dissolved gas abundances are affected by processes deep within the hydrothermal system. Fluid samples collected from research wells, subject to the least contamination by atmospheric gases, were found to be markedly depleted in noble gases compared with the expected values for 20°C air-saturated water. Some samples carried a radiogenic component, as indicated by 4°Arp6Ar ratios higher than the atmospheric value and by strong enrichments in 4He relative to the other noble gases (Mazor &
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YNP HYDROTHERMAL SYSTEM 33 Fournier 1 973). Craig et al ( 1978) and Welhan ( 198 1 ) presented 3HerHe isotope data that indicate a subcrustal component of gas eHe) in the Yellowstone thermal waters. The largest leakage of 3He was detected at Mud Volcano, where a 3Hej4He ratio about 1 6 times the air ratio was detected. More extensive sampling by Kennedy et al (1985) confirmed the results of Craig et al (1978) and Welhan (1981) and provided broader coverage throughout the Park.
Kennedy et al ( 1985) analyzed the abundances of five stable noble gases and the isotopic compositions of He, Ne, and Ar for 66 gas and water samples representing 1 9 different regions of hydrothermal activity. The isotopes of Kr and Xe were also determined in a few samples and found to be of normal atmospheric relative abundance. Kennedy et al explained their results by mixing of three source components: magmatic, crustal, and atmospheric, each having a characteristic noble gas isotopic range, with high 3He indicative of a magmatic source. In addition to the 3He anomaly at Mud Volcano, Kennedy et al ( 1985) found 3Hej4He 1 3.4 times the air ratio at Gibbon, 1 0.4 times the air ratio at Crater Hills, 9.5 times the air ratio at Shoshone, and generally about 7 times the air ratio elsewhere within the caldera. Except in the southwest region of the Park at Boundary Creek, the maximum 3Hej4He decreased rapidly to values < 3 times the air value immediately outside the caldera. High 3Hej4He anomalies were not found associated with either of the two zones of low seismic velocity (4.0 km S- I) reported by Lehman et al ( 1 982) that have been widely cited as possibly indicating a body of partially molten material at a depth of 3-4 km beneath the caldera. The lack of 3HerHe anomalies associated with the western anomaly is not surprising in view of more recent data that have shown that the western anomaly is fictive (Brokaw 1 985). The anomaly at Mud Volcano and Crater Hills was attributed to a relatively shallow depth to the Curie-point isotherm and coincidence with a zone of maximum present-day uplift within the caldera (Kennedy et al 1 985). The 3Hej4He anomaly at Gibbon was considered to be somewhat of an enigma, possibly reflecting the fact that the youngest rhyolite lavas in the Park (Doe et al 1 982) occur there.
Kennedy et al ( 1 987) studied 1 2 different springs in Lower Geyser Basin and found that the 3HerHe ratio ranged from � 2.7 to 7.7 times the air ratio and increased with total bicarbonate concentration. Because the total bicarbonate concentration is a function of the availability of CO2, which in turn is a function of deep boiling with phase separation prior to COT bicarbonate conversion (as previously discussed), Kennedy et al ( 1987) concluded that boiling and dilution of a single thermal fluid could explain their gas data for Lower Basin. They also noted that a fluid that boils to a greater extent loses a larger proportion of 3He, as well as CO2, leaving
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34 FOURNIER
a helium-poor residual fluid in which the isotopic composition of helium will be strongly affected by the addition of radiogenic helium.
An alternative interpretation of the district-wide helium data of Kennedy et al ( 1985), based on the Lower Basin model of Kennedy et al ( 1987), is that 3HetHe ratios throughout the caldera and in immediately adjacent regions are controlled mainly by boiling and mixing. Within most of the western part of the 0.6-Ma Yellowstone caldera, where locally derived meteoric waters circulate relatively freely through rhyolite lava flows, a considerable amount of mixing of the deep component with shallow water occurs, followed by adsorption of radiogenic 4He from the surrounding permeable rocks. Thus, an initially high 3HerHe ratio that is about 1 6 times the air ratio everywhere deep within the caldera may be generally reduced to less than 7 times the air ratio at shallow levels. Exceptions are in those few places where little shallow mixing of the ascending liquid occurs, and where steam that separates at very high temperatures makes its way to the surface with little or no dilution by other steam formed at shallow levels where local meteoric water boils. It may be more than coincidence that the highest 3HerHe ratios all have been found at the margin of or outside the boundary of the 2.0-Ma caldera shown by Christiansen ( 1 984) and, therefore, outside the region where relatively permeable rhyolite flows are likely to be found deep within the caldera between different ash-flow tuff units, as discussed previously.
At Crater Hills the one flowing hot spring (Table 1 , number 1 2) is located within the 0.6-Ma caldera and at the projected eastern edge of the 2.0-Ma caldera. This is a likely site for prolonged and intense hydrothermal activity since the formation of the first caldera. The hot spring at Crater Hills has the highest chloride concentration of any thermal water in the Park ( 1044 mg kg- 1 measured by Allen & Day 1 935) and the highest chemical geothermometer temperatures (270-300°C). This water appears to represent the most direct upward leakage of fluid from deep in the system.
Enthalpy-chloride relations (Figure 7) suggest that the parent water deep in the system (point R) has a minimum enthalpy of about 1 700 J g- l (about 350°C) and contains about 450 mg kg- 1 Cl. Water cools adiabatically by decompressional boiling in moving [rom point R to the reservoir feeding the hot spring at Crater Hills, but evidently the evolved steam and other gases remain essentially in contact with the residual liquid. This requires relatively impermeable surrounding rock. The proximity to the Mud Volcano vapor-dominated system suggests that this is likely to be the general situation in that part of the Park because low permeability is required to maintain a hydrologically underpressured zone next to a zone of hot-water upflow from deep in the system. Low permeability within
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YNP HYDROTHERMAL SYSTEM 35
the upper part of the stratigraphic section also provides a mechanism by which steam that forms deep under the Mud Volcano region rises to the surface with little contamination by gases in steam derived by boiling, shallow, local meteoric water. At Gibbon the hot-spring waters and gases issue from relatively impermeable ash-flow tuffs. There the enthalpy-chloride relations (Figure 7) show that the water in the local shallow reservoir has about the same chloride concentration as that in the deep parent water. Therefore, the Gibbon water may have cooled conductively without mixing while flowing from deep in the system to a more shallow level. In this event, the initial concentrations of noble gases in the deep water would tend to be preserved, modified mainly by leaching of radiogenic components from the shallow reservoir rocks.
METEORIC RECHARGE OF THE HYDROTHERMAL SYSTEM
Craig et al ( 1 956) compared stable isotopes in Yellowstone thermal waters with those in local meteoric water and concluded that any possible magmatic water contribution is less than a few percent. Although the interpretation of the stable isotope compositions of hot-spring waters is made difficult by considerable dilution of ascending deep water by relatively locally derived meteoric water and by subsurface boiling that results in fractionation of isotopes between steam and the residual liquid, the general conclusions of Craig et al ( 1 956) still appear to remain valid after the collection and analyses of many more samples.
Taking account of variations in chloride concentrations and of oxygen and deuterium isotopic abundances in thermal waters in the Park that result from boiling and mixing, Truesdell et al ( 1 977) proposed two endmember models for the behavior of steam that forms by decompressional boiling of an ascending solution. The first is a '''single-stage'' model in which all the steam remains mixed with the water and separates near a single temperature, and the second is a "continuous" model in which steam is physically separated from the water as it is formed. A multiple-stage model that produces results intermediate between single-stage and continuous steam separation appears to be applicable to many natural hotspring waters in Yellowstone. Truesdell et al ( 1977) assumed a temperature (360°C) and chloride cvncentration (3 10 mg kg- I ) of a deep thermal water (taken from the enthalpy-chloride model of Truesdell & Fournier 1976) and used the measured temperature, chloride concentration, and isotopic composition of water collected from the Y- 1 2 drill hole to calculate the relative deuterium concentration «(jd) of the deep-water component, using the (j (del) notation in which the isotopic concentration is reported as parts
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36 FOURNIER
per thousand difference from a standard, where
(jd = [(rsample/rsMow) - l] x 1 03, ( 1 )
rsample i s the ratio of deuterium to hydrogen in the sample, and rSMOW is the ratio of deuterium to hydrogen in standard mean ocean water. They concluded that the chloride-(jd relations for Norris, Lower, Upper, and Shoshone Geyser Basins are compatible with boiling of mixtures of the deep water with cold, dilute (SOC, 2 mg kg- I Cl) meteoric water, varying in (jd according to locality in the Park. In addition, the data strongly suggest that recharge to the deep thermal water must be either precipitation from areas remote from the geyser basins or ancient precipitation from a cooler period. The major conclusions of Truesdell et al ( 1977) would not be changed significantly by assuming a higher temperature for the Y-1 2 water (it probably cooled conductively from at least 270°C rather than 238°C) and a higher chloride concentration in the deep, 360°C thermal water. Unpublished stable isotope analyses of an extensive suite of samples collected and analyzed by Robert O. Rye and additional stable isotope analyses by Tyler B. Coplen indicate that the most likely recharge area for the deep component of the geothermal water is in mountainous regions north and northwest of the caldera, where deep snow accumulates each winter.
GEOTHERMAL HEAT FLUX
The total convective heat flux in chloride-rich water and associated steam from all the Yellowstone thermal areas east of the Continental Divide was estimated in 1 966-67 to be about 4.0 x 1 0 1 6 cal yr- I , or 5 .3 x 1 09 W by the river chloride inventory method (Fournier et al 1 976). Nearly all the chloride discharged by the hot-spring waters eventually enters the four major rivers that drain the Park: the Yellowstone and Madison Rivers east of the Continental Divide, and the Snake and Falls Rivers west of the Divide (Figure 1 ). The Firehole River (draining the Upper, Midway, and Lower Geyser Basins) and the Gibbon River (draining Norris Geyser Basin) join to produce the Madison River. Fournier et al ( 1976) used enthalpy-chloride information (Figure 7) to estimate the chloride concentrations of the waters in the shallow reservoirs feeding individual geyser basins and the chloride concentration of the water deep in the hydrothermal system before decompressional boiling and before near-surface mixing (point P with 400 mg kg- I Cl). Measured rates of discharge of the rivers and of the chloride carried by those river waters were then used to calculate the rate of total discharge of the deep component of the hotspring water. The method automatically compensates for boiling and mixing of thermal waters as they rise. Corrections are required, however,
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YNP HYDROTHERMAL SYSTEM 37
for background chloride contributed to the rivers by precipitated rain and snow and by low-temperature weathering of rock. The background Cl generally amounts to about 1-2 mg kg- l in the non thermal component of the river water, depending on the proportion of water from different non thermal sources. (For example, snow melt that has not penetrated into the ground is lower in Cl than nonthermal groundwater that has been underground a long time in contact with various types of rock). A summation of the results for the Yellowstone and Madison Rivers, plus an additional 5% assumed for the thermal discharge into the Snake and Falls Rivers, gave a figure of 10 1 1 kg yr- l for the deep component of the hydrothermal water and 4 x 107 kg yr-l for the chloride carried by that water in 1966-67 (Fournier et al 1 976, Fournier & Pitt 1985). For a minimum initial water temperature of 340°C, the total amount of heat discharged by that water is 5 .3 x 109 W, which translates to an average heat flux of about 2000 mW m-2 over the entire 2500 km2 of the Yellowstone caldera. The above estimates are not changed significantly when the higher chloride and higher enthalpy values of point R in Figure 7 are used for the deep water instead of those of point P.
Norton & Friedman (1985) reported results of monitoring the chloride flux from the Yellowstone, Madison, Snake, and Falls Rivers for the 1 983 water year, extending from September 1, 1982, through August 3 1 , 1 983. They calculated a total chloride flux of about 5.8 x 107 kg yr- 1 in these river waters, of which about 5 .5 x 107 kg yr- l were estimated to be contributed by thermal water. They also found that the chloride flux from thermal waters into the drainage west of the Continental Divide is about 26% of the total for the Park, much higher than the 5% estimated by Fournier et al (1 976). When 26% thermal water, rather than 5%, for the region west of the Continental Divide is applied to the 1966-67 data, a revised flux of thermal water chloride of 4.8 x 1 07 kg yr- 1 is obtained. This is 1 7% less than the thermal water Cl flux for the Park reported by Norton & Friedman ( 1985).
The chloride flux measured in 1983 appears to be exceptionally high compared not only with the 1966-67 measurements, but also with measurements made in subsequent years (USGS, unpublished data). Figure 8 shows calculated instantaneous thermal water chloride fluxes for the Madison River during the 1 983 and 1 985 water years and during 1 966-67. A correction of 1 mg kg- 1 Cl for the non thermal contribution has been applied to the data shown in Figure 8. The wide scatter in the data after day 250 corresponds with periods of high river discharge during the spring and early summer melting of snow. During high stages of river discharge, the calculated fluxes of chloride have a greater uncertainty because small percentage errors in chloride analyses and flow measurements are magnified and because discharge measurements, obtained from rating curves,
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1 500
U Ql
� 1 000 0'
X :::J '+-
Ql -0 500 'C 0 ..c U
e 0... I
Madison River
O �rr .... -rrr""rrTT .. -rrr .. -rrrTT,,-r .... -r o 50 1 00 1 50 200 250 300
Day number (water year) 350
Figure 8 Calculated chloride flux vs. day of water year for the Madison River. Water year starts September 1 and ends August 3 1 of succeeding calendar year. Triangles indicate data for 1966, dots for 1 967, circles for 1 983, and asterisks for 1985.
are less precise as a result of changing channel profiles that require recalibration of rating curves. Also, during high stages of river flow, slight differences in the assumed background chloride correction make a big difference in the calculated flux of thermal water, However, uncertainties in the background chloride correction can be minimized with the aid of a log-log plot of river discharge rate versus river chloride. This is done by extrapolating to the assumed CI concentration in the deep parent water using theoretical mixing curves that have fixed shapes, as shown in Figure 9. By this method it is easy to see at what discharge rate the correction for background chloride becomes negligible.
Using data in Fournier et al ( 1 976) and unpublished data of the USGS, I have calculated approximate ranges of the instantaneous flux of total thermal water discharged by all the chloride-rich hot-spring systems in the Park for the years 1 966-67 and the water years 1 983-86 (Figure 1 0) . In assembling the data with which Figure 10 was constructed, no attempt was made to check the validity of the most extreme data points or to use conductivity data to estimate chloride concentrations between times when water samples were collected for analyses. The circles show where most of the data plot for low stages of river flow. The data presented in Figure 1 0
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Cl .., ..... Cl E
1 00
1 0
.. 400
+ 1966 -67 • 1984-85
YNP HYDROTHERMAL SYSTEM 39
1 �������������������� 1 1 0 100 1 000
Flow, cfs
Figure 9 Logarithmic plot of flow rate vs. dissolved chloride for the Gibbon and Firehole Rivers. Calculated mixing curves are for an assumed end-member deep thermal water containing 400 mg kg- 1 chloride and non thermal waters containing I and 2 mg kg- 1 chloride, respectively.
6
5
4 " QJ �
co 3 E t t � 0 u:: 2
O �--�--��----L---�--�--� 1 966-67 1 983 1 984 1 985 1 986 Water year
Figure 10 Range in calculated instantaneous fluxes of thermal water vs. year. Circles show most common flux during low stages of river flow.
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clearly show about a 1 5-20% increase in flux of thermal water and heat during the 1983 water year compared with the other years. Interestingly, this anomalously high flux of thermal water occurred during the year preceding the 7.3 magnitude M t. Borah earthquake on October 28, 1 983. A little over three months before that earthquake the flux of thermal water decreased to what appeared to be more normal values, and then a month before that earthquake the flux of thermal water again increased to anomalously large values. Apparently, Norton & Friedman ( 1985) did not directly compare their data with the results of measurements made in 1966-67 at comparable stages of river flow and did not conclude that there were significant differences between the two data sets, except in regard to the assumed flux of chloride from thermal water into the Snake and Falls Rivers. However, they did note an increase in chloride flux commencing early in October 1 983 and culminating shortly after the Mt. Borah earthquake. Future detailed analyses of river flow, temperature, and conductivity data collected hourly from 1 982 through 1 987 are likely to reveal significant variations in the total flux of thermal waters since 1983.
EFFECTS OF GLACIAL ICE COVER ON THE HYDROTHERMAL SYSTEM
A series of research wells drilled at several thermal localities in Yellowstone Park (White et al 1 975) have provided information about present temperature-pressure-depth relations. Near major zones of active convective upflow of thermal water, present temperatures follow boiling-point curves appropriate for water tables at the present land surface, with hydrostatic pressures equivalent to a cold rather than hot column of overlying water (Fournier 1983). Muffler et al ( 1971 ) postulated that boiling-point curves and rock temperatures were raised within the upper parts of some hotspring systems at Yellowstone during times of glacial cover, and that sudden decreases in the water table (as by draining of an ice-dammed lake) resulted in major hydrothermal explosive activity. Primary and secondary fluid inclusions in material from several cored holes in Yellowstone National Park (White et a1 1975) have been studied by Bargar & Fournier ( 1 988) to quantify thermal effects that may have accompanied a raising of the water table and the corresponding boiling-point curve during times of thick glacial ice cover. The results from two holes (Y-6 and Y- 1 3) have been reported as parts of other investigations (Bargar & Beeson 1 984, Bargar et al 1 985).
Mineral samples containing fluid inclusions suitable for study were found in drill cores from 8 of 1 3 wells drilled by the USGS. Primary, pscudosecondary, and secondary liquid-rich fluid inclusions in quartz and fluorite were utilized to determine a range of homogenization temperatures
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(Th) a t given depths. Within zones o f major hydrothermal upflow, fluid inclusion Th values all were found to be equal to or higher (commonly 20-50°C and up to 1 55°C higher) than present temperatures at the depth of collection. The greatest range in Th values was found in material collected from the Y-I 3 drill hole in Lower Geyser Basin (Bargar et al I985). In the absence of glacial ice, topographic constraints place a limit of about 1 20-1 50 m on the maximum depth of any lake occupying Lower Geyser Basin. The elevation of the water table over Lower Geyser Basin required to raise boiling-point curves sufficiently to account for the observed fluid inclusion Th values ranges from 425 m (a conservative estimate that does not include the highest Th values because of possible necking-down phenomena after entrapment) to 750 m (a liberal estimate that includes all the data). Thus, it appears that glacial ice is required to account for the boiling-point curve elevation. The estimated thickness of the glacial ice over Lower Basin during the last glaciation required to explain the fluid inclusion data is within the range estimated by other criteria (K. L. Pierce, written communication, 1988).
Away from zones of major upflow, Th values were all found to be equal to or lower than present temperatures (commonly 1 0-20°C and as much as 60°C lower). The lower paleotemperatures at the margins of hydrothermal upflow may be attributed to lower rates of discharge of warm water and/or higher rates of recharge of cold water in the past, probably during periods of glacial cover.
INTERACTION OF CIRCULATING WATER WITH THE HEAT SOURCE
In most current models of magmatic-hydrothermal systems (Lister 1 974, 1980, 1 983, Hardee 1 982, Sleep 1 983, Fournier & Pitt 1985), relatively dilute meteoric water circulates along fractures above and to the sides of a magmatic heat source. The water is separated from very hot crystalline rock or magma by an impermeable zone where quasi-plastic deformation prevents brittle fracture and maintains low permeability. Heat extracted by the circulating water slowly cools the system, allowing thermal cracking and faulting to progress inward so that water steadily migrates deeper into the heat source while attaining about the same maximum temperature. Although there is no direct evidence about the maximum temperature attained by the convecting hydrothermal fluid, enthalpy-chloride relations (Figure 7) provide a means of estimating a likely minimum temperature deep in the system, about 340-350°C. An approximate upper limit for the temperature likely to be attained by the water in the hydrothermal system that convects to the surface can be estimated from positions of boilingpoint and condensation-point curves (Fournier 1 987) in temperature-pres-
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42 FOURNIER
sure space for a dilute chloride-rich water about equivalent in salinity to a 0.7 wt% NaCI solution (Fournier & Pitt 1985). For the restricted range of compositions of the relatively dilute solutions that issue at the surface at Yellowstone, these fluids could not have attained temperatures higher than about 420°C at a depth of 4 km and 430°C at 5 km. Exceeding these temperatures would have resulted in an ascending solution intersecting the stability field of gas plus brine, resulting in the formation of a small amount of highly saline brine and a large amount of very dilute steam (Fournier & Pitt 1985, Fournier 1987).
The likely depth of circulation of hydrothermal fluids also is a matter of conjecture. Deep drilling in other presently active hydrothermal systems has shown that permeability at temperatures above about 300°C is controlled mainly by fractures. Precipitation of minerals and quasi-plastic flow are particularly efficient at decreasing permeability by closing fractures at temperatures above about 350-400°C (Fournier & Pitt 1 985). Therefore, sustained circulation of hydrothermal fluids at these high temperatures is likely to take place only where the rocks are subjected to repeated brittle fracture that counteracts the processes that decrease permeability. Focal depths of well-located earthquakes are deeper than 1 5 km outside the Yellowstone caldera, but within the caldera they seldom exceed 4-5 km. Evidently the rocks below 3-4 km are too weak as a result of high temperatures (> 350°C) to sustain sufficient differential stress for seismic activity to occur (Smith & Bruhn 1984, Fournier & Pitt 1 985). Therefore, the maximum depth of sustained fluid flow at hydrostatic pressure is probably about 4-5 km within the caldera and deeper at the sides.
If the assumption is made that thermal energy is transferred uniformly by conduction from the heat source across a 2500 km2 surface, through hot dry rock to the base of circulation of the hydrothermal system, a thermal gradient of about 700° to l OOO°C km - I is required to sustain the average convective thermal output (Morgan et al 1 977, Fournier & Pitt 1 985). To account for both the large convective heat flux and the lack of earthquakes below about 4-5 km, it is likely that magmatic temperatures are attained beneath some portions of the caldera at depths as shallow as about 4.5-5.5 km, about the same depth range indicated by the various other geophysical methods.
CRYSTALLIZATION OF MAGMA AS A SOURCE OF HEAT AND AQUEOUS FLUIDS
If the heat carried in the advective flow of water from the system were supplied entirely by the latent heat of crystallization of rhyolite magma,
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about 5 x 10" kg yr- 1 of magma (about 0.2 km3 yr- 1 ) would be required to furnish that heat (Fournier & Pitt 1 985). In contrast, cooling already crystalline rock by 300°C (to reduce the temperature from about 730° to 430°C) will furnish about the same amount of heat that can be obtained from the latent heat of crystallization of the same volume of rhyolitic magma (using a latent heat of crystallization of75 cal g- l and heat capacity of 0.25 cal g- l °C- 1) . Thus, crystallizing about 0 . 1 km3 of rhyolitic magma and then cooling it about 300°C could furnish the heat carried to the surface by the hot-spring system. A lack of water-bearing minerals in the rhyolite erupted at Yellowstone coupled with the presence of the mineral fayalite suggests that the initial magma contained not more than 1 to 2 wt % dissolved water (Hildreth et al 1 984). The chloride contained in fresh obsidian glass shows that the minimum initial chloride concentration in the magma was 1 500 mg kg-1 (Wes Hildreth, oral communication, 1 984). Strong partitioning of this chloride into an aqueous magmatic fluid that separates upon slow crystallization of the magma will produce a fluid containing a minimum of 1 2 to 25 wt% salt. The maximum quantity of water that could be liberated by crystallizing 0.2 km3 yr- I of magma initially containing 2 wt% water is 8 .8 x 1 09 kg yr- 1 , and a minimum of about 7 x 1 08 kg yr- 1 of chloride would be liberated with that water. For comparison, recall that about 10" kg yr- 1 water (not including water incorporated by near-surface dilution) and 4 x 1 07 kg yr- 1 chloride are discharged by the hot-spring system. Thus, a maximum of about 9 wt% of the deep component of hot-spring water could be magmatic, but only about 6 wt% of the available chloride can be discharged with that water. The deficiency of chloride in the discharged thermal water may be interpreted in two ways: (a) Thermal energy is supplied mainly by cooling already crystalline rock with little thermal input from the latent heat of crystallization of magma; or (b) brine evolved from presently crystallizing magma is ponding beneath the dilute hydrothermal system, and little of this magmatic water is incorporated into the convecting meteoric system. (A mixture of about 0.2-0.4% magmatic fluid with 99.6-99.8% meteoric water could account for all the chloride discharged by the Yellowstone hot springs.) It was noted previously that stable isotope data also suggest that very little, if any, magmatic water is discharged by the hot springs (Craig et al 1 956, Truesdell et al 1 977). The preferred model of Fournier & Pitt ( 1985) is one in which the thermal energy carried to the surface by the convecting hot-spring water comes in part from crystallizing magma and in part from cooling crystalline rock, and that relatively dense brine liberated from the crystallizing magma is accumulating at depth beneath dilute, much less dense hydrothermal waters.
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LONGEVITY OF PRESENT HYDROTHERMAL ACTIVITY
Uranium series disequilibrium studies (Sturchio et al 1 987) of cores from the Y-7 drill hole in Biscuit Basin, Yellowstone National Park, indicate that there was significant uranium mobility 1 5,000-20,000 yr ago, corresponding to a period of time when glaciation would have allowed temperatures to be significantly higher at the depths penetrated by the drill hole. Unpublished uranium series disequilibrium studies of cores from other drill holes show that vein ages are generally much younger than the ages of the volcanic host rocks (N. C. Sturchio, written communication, 1 988).
Bargar & Fournier ( 1 988) observed that Th values of primary and secondary liquid-rich fluid inclusions trapped in hydrothermal minerals sampled from within major zones of up flow are all equal to or greater than presently measured rock temperatures at the point of collection. This is strong evidence that the thermal waters have not been cooler than present temperatures since formation of the host minerals. Also, because the presence of glacial ice seems to be required to explain elevated boilingpoint curves, it can be concluded that hydrothermal activity beneath the major geyser basins has been continuous since sometime within or before the last glaciation, a minimum of 1 4,000-45,000 yr. The fluorite and much of the hydrothermal quartz used in the Bargar & Fournier ( 1 988) study were deposited relatively late in the sequence of hydrothermal events, after most of the surrounding rock had already been thoroughly altered by hydrothermal activity, but before the last glacial ice decreased in thickness to less than 450-750 m. The data are consistent with a model in which hydrothermal activity within the 0.6-Ma Yellowstone caldera started soon after the formation of that structure and continued unabated thereafter, with significant variations in intensity and places of discharge in response to pulses of volcanism and the waxing and waning of glacial episodes.
Ovcr thc minimum time span during which hydrothermal activity appears to have operated at about its present intensity ( 1 5,000 yr since the end of the last glaciation), a largc amount of thermal energy has been extracted from the system, both by convection and conduction, as indicated above. Ifwe use the assumed thermodynamic properties of materials given in Fournier & Pitt ( 1 985), the calculated amount of heat extracted just by advection of chloride-rich water from the hydrothermal system in 1 5,000 yr could be supplied by subsolidus cooling over a 350°C range (for example, from 750° to 400°C) of a crystalline body of rock 1 .2 km thick beneath the entire caldera (2500 km2), or by the latent heat associated with crystallizing a body of magma 1 .2 km thick beneath the entire caldera with
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no change in temperature of the system. Crystallizing a body of silicic magma 0.6 km thick and then cooling it by 350°C would also provide the required thermal energy. Again, these are minimum requirements, as the total flux of heat from the system is undoubtedly larger than the number estimated from the chloride flux. The chloride flux method does not provide information about hear' carried to the surface by steam separated from convecting liquid that cools and flows back downward into the system, or about a purely conductive component of heat flow.
VOLUMETRIC CHANGES ASSOCIATED WITH CRYSTALLIZATION OF MAGMA
Cooling of hot but already crystalline rock should lead to thermal contraction and a subsidence of the caldera floor. However, a significant uplift of the caldera has been measured over the period 1 923-84 (Pelton & Smith 1 979, 1 982, Dzurisin et a1 1 986, Dzurisin & Yamashita 1 987). This inflation stopped in 1 985, and deflation occurred in 1986 (Dzurisin & Yamashita 1 987). Studies of raised and submerged shoreline terraces show that there has been episodic uplift interspersed with local deflation of the caldera throughout the Holocene (Hamilton 1 985, Meyer & Lock 1 986). Meertens & Levine ( 1985) have suggested that the uplift is the result of horizontal compressive strain, but this does not appear to agree with focal mechanisms of earthquakes within the caldera and the generally extensional tectonic environment of the region. Also, the average horizontal strain rates measured in the northeast part of the caldera and in the Hebgen Lakc region to the northwcst of thc caldcra from September 1 984 to August 1 987 appear to preclude regional tectonic movements as a primary cause of the inflation and deflation (unpublished data of J. C. Savage and D. Dzurisin).
Two mechanisms that might account for both the episodic inflation and deflation of the caldera each require crystallization of rhyolitic magma in amounts that are compatible with the thermal budget for the magmatichydrothermal system. Both mechanisms may operate to varying degrees at the same time. The first involves episodic intrusion of basalt from the mantle into the upper crust at the base of a silicic magma chamber at a depth of about 1 0- 1 5 km, and continuous crystallization of rhyolitic magma at a depth of about 4-5 km. The thermal energy discharged by the advecting thermal waters at Yellowstone could be supplied by the latent heat of crystallization of about 0.2 km3/yr of rhyolitic magma, with no change in temperature within the igneous part of the system (Fournier & Pitt 1 985). If purely conductive heat loss and heat carried upward by convecting hydrothermal fluids that do not discharge at the surface also
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were taken into account, the amount of magma required to supply the thermal flux from the magmatic-hydrothermal system would be correspondingly higher. Crystallization of 0.2 km3 of rhyolitic magma would lead to a volumetric contraction of about 0.0 1 4 km\ provided that the evolved, aqueous-rich magmatic fluid all escaped into the overlying hydrostatically pressured hydrothermal system. Thus, if the latent heat of crystallization of magma is an important source of thermal energy for the hydrothermal system and if the evolved magmatic fluids continuously and efficiently escape from the system, then the normal state of deformation of the caldera is likely to be deflation at a rate that is close to the observed rate of subsidence since 1 985. Inflation will occur only during periods of episodic intrusion of basalt from the mantle into the crust at a rate fast enough to more than compensate for the decrease in volume resulting from crystallizing rhyolite magma. Stability at the caldera floor will occur only when the volume increase resulting from basalt intrusion at the base of the relatively high-level magma chamber just balances the volume decrease resulting from crystallizing rhyolite at the top of the chamber
The second mechanism that can account for rapid fluctuations from inflation to deflation involves entrapment of aqueous magmatic fluids at lithostatic pressure beneath self-sealed rock and the episodic release of these fluids into the hydrostatically pressured system (Fournier & Pitt 1 985, Fournier 1988). Self-sealing at the base of the hydrostatically pressured hydrothermal system may occur as a result of chemical transport and precipitation of vein minerals and ductile flow of rock that becomes relatively weak at temperatures exceeding 400°C (Fournier & Pitt 1 985). lf the magmatic fluid that is liberated during crystallization of 0.2 km3 yr- I of rhyolitic magma containing about 2 wt% water is trapped at litho static pressure at a depth of 4-5 km, there will be a net volume increase of about 0.026 km3 yr- I ; more than enough to account for the observed rate of inflation of the caldera from 1 923 to 1 984. The density of the magmatic gas phase is not particularly sensitive to the initial ratio of water to chloride in the magma, because at the likely depth and temperature of crystallization, the magmatic fluid will dissociate to a small amount of very saline brine plus relatively dilute gas, as previously discussed. Episodic hydro fracturing and injection of pore fluids that have accumulated at lithostatic pressure into the hydrostatically pressured hydrothermal system would result in episodic deflation of the caldera. The pressure surge associated with this injection of magmatic fluid may be accommodated entirely or in part by an increased rate of flow of thermal water from the system, by a decrease in the rate of recharge of cold meteoric water into the system, or by condensation of a relatively small amount of steam accompanied by a slight rise in the level of the interface between liquid and steam within a
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deep vapor-dominated region that appears to be widespread in the eastern part of the Park. Also, episodic rupturing of self-sealed rock that separates regions of litho static and hydrostatic pore-fluid pressures could account for some of the seismic swarms that occur within the caldera.
Although the above calculations and speculations do not demonstrate conclusively a causal relationship between the uplift and liberation of fluid from crystallizing magma, they do show that the volumetric effects of liberating a magmatic fluid could be important if the latent heat of crystallization of magma is an important source of energy for the hydrothermal system. Future attempts will be made to determine whether a statistically significant increase in the rate of hot-water discharge has coincided with a change from inflation to deflation of the caldera. A major problem in the interpretation of the data is the large effect that the Mt. Borah earthquake appears to have had upon the hot-spring discharge rate. However, even if there is no detectable increase in the rate of discharge of thermal water coinciding with the onset of deflation, we cannot rule out the possibility that depressuring of magmatic water is a major contributing factor, because the volumetric effects may be most pronounced within the cold recharge part of the system and within the vapor-dominated zones.
MANTLE BASALT AS THE ULTIMATE SOURCE OF THERMAL ENERGY AT YELLOWSTONE
It is generally agreed that continuing injection of basaltic magma ( 1 200-I 300°C) from the mantle into the crust over a long period of time supplies the thermal energy to melt crustal material, forming large volumes of silicic magma at 800-900°C (Christiansen 1 984). Crystallizing and cooling ( 1 200° to 860°C) about 0.08 km3 yr- I of basalt could furnish the heat for the present-day Yellowstone hydrothermal system. This yearly volume of basalt is about seven times more than the historic measured average rate of volumetric inflation of the Yellowstone caldera. If mass (either cooled basaltic magma or dense crystalline material) is not convecting or settling downward into the lower crust and mantle beneath the Yellowstone caldera to counterbalance the buoyant upflow of hotter basaltic magma, then the Yellowstone volcanic system as a whole is cooling at present and the amount of silicic melt (if any) in the shallow part of the system is steadily decreasing.
CONCLUSIONS
Meteoric water, derived in great part from precipitation on mountains to the north and northwest of the Yellowstone caldera, convects at hydro-
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48 FOURNIER
static pressure to a depth of about 4--5 km within the caldera and attains a maximum temperature of about 350-430°C. This water picks up about 300-500 mg kg- I Cl from the rock through which it passes and also picks up gaseous components liberated from magma and crystalline magmatic rocks. Cation ratios are fixed by temperature-dependent water-rock reactions. There is a possibility that some or even most of the chloride in the relatively dilute hydrothermal system is derived by interaction with an underlying brine of magmatic origin that is postulated to be present at depth. Even though some variation in chloride is likely to occur in the deepest parts of the convecting system, these variations appear to be small, probably because rocks with similar compositions occur at depth throughout the caldera, and because there has been a lot of flushing of dissolved constituents by hydrothermal advective flow throughout a 2-Myr interval of volcanic activity. Thus, the assumption in the application of boiling and mixing models that a deep "parent" water has a uniform chloride concentration is probably a good first approximation for use in estimating conditions deep in the system.
Heated fluid convects upward and to various degrees in different places boils and/or mixes with shallow fluids of more local meteoric origin. Waterrock chemical reequilibration occurs in relatively shallow local reservoirs that directly feed individual geyser basins. Variations in dissolved gases in general, and in 3HerHe ratios in particular, can be explained entirely in terms of variation in underground boiling and mixing of deep and shallow waters. During periods of glaciation, temperatures in zones of major hydrothermal upflow increased at given depths as a result of elevation of boiling-point curves by increased hydrostatic head. Vapor-dominated systems in the eastern part of the Park formed at the close of the last glaciation partly as a result of a lowering of the water table by flow toward the Grand Canyon of the Yellowstone and partly in response to a large input of thermal energy beneath a thick sequence of relatively impermeable rock. The convective heat discharged by the hydrothermal system is generally about 4 x 10 1 6 cal yr- 1 , but an approximate 20% increase was observed in 1 983, the year preceding the Mt. Borah earthquake. Hydrothermal activity has been continuous at about its present level for at least 1 5,000 yr and probably much longer. It is not known whether the magmatic-hydrothermal system, as a whole, is presently cooling down or heating up. The thermal energy that drives the hydrothermal convection can come entirely from the cooling of already crystalline rocks, entirely from the latent heat of crystallization of silicic magma at a depth of about 4--5 km, or from a combination of the two. The recent uplift of the Yellowstone caldera suggests that the injection of magma into the system and/or the liberation of fluid from a crystallizing magma are processes
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that are presently active. If magmatic fluids are presently being liberated from a crystallizing siliceous magma at a depth of 4-5 km, they should dissociate to highly saline brine plus relatively dilute gas. Very little (if any) of that brine is becoming incorporated into the portion of the hydrothennal system that convects to the surface. The magmatic fluid is ponding beneath the dilute fluid and/or is being trapped in a region where pore-fluid pressures are lithostatic. Accumulation of magmatic fluids at litho static pressure and episodic injection of portions of this fluid into the hydrostatically pressured system may account for the observed inflation and deflation of the caldera. However, the observed uplift and subsidence also may be accounted for by episodic intrusion of basalt from the mantle into the upper crust at a depth of about 1 0- 1 5 km and continuous crystallization of silicic magma at about 4-6 km without accumulation of magmatic gases at lithostatic pressure. Uplift would occur only when the rate of basalt intrusion more than compensated for the normal subsidence resulting from crystallization of rhyolitic magma. At this time the most plausible model for the magmatic-hydrothermal system is one in which the thermal energy carried to the surface by the advecting hot-spring water comes in part from crystallizing silicic magma and in part from cooling crystalline rock, and in which brines are accumulating deep in the system. It is likely that crystallization and cooling are occurring at some levels while new magma accumulates at others.
ACKNOWLEDGMENTS
Since I first started working on the Yellowstone hydrothermal system in 1 969, I have benefited greatly from collaboration with many colleagues in the USGS, with special thanks to A. H. Truesdell, L. J. P . Muffier, J. M. Thompson, G. W. Morey, J. J. Rowe, and particularly Donald E. White, who first introduced me to the joys of studying the hydrogeochemistry of Ycllowstone's thermal waters and who has tried to keep me on the right track ever since. Over the years the National Park Service (NPS) has been and continues to be very supportive of our research in the Park. R. A. Hutchinson of the NPS has been especially helpful in collecting hot-spring water samples for analyses from remote regions and in monitoring the chemical behavior of various thermal waters over many years. D. R. Norton and Irving Friedman of the USGS have been deeply involved with a cooperative USGS-NPS program to monitor the chloride flux in the major rivers in Yellowstone since 1982, and they have provided me with much of the data for calculating the recent convective heat flux from the system.
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