Southeastern Idaho spring/well waters
Compositions of waters from hot/warm springs and wells in southeastern Idaho
are presented in Table 2. The pH values of southeastern Idaho thermal waters are
circum-neutral, ranging from 6.2 to 8.1, with arithmetic mean, median, and
standard deviation of 6.87, 6.70, and 0.51, respectively; the field
temperatures range between 20 and 84 ∘C. The aqueous
chemistry of these thermal waters shows a large range in total dissolved
solids (TDS) from about 250 mg L-1 (Downata Hot Spring) to more than
14 000 mg L-1 (E. Bingham Well).
The dominant cations in the southeastern Idaho thermal waters are Na and Ca
with minor amounts of Mg (Fig. 2). The thermal waters include samples
dominated by Cl, HCO3, or SO4 while others appear to be
dominated by more than one anion. Hierarchical cluster analysis using Ward's (1963)
method as implemented in SYSTAT 13 (SYSTAT Software, Inc.) was
performed using the 6 Piper diagram (Fig. 2) end-members (Ca2+,
Mg2+, Na++K+, Cl-, HCO3-+CO32-,
SO42-) for classifying water compositions. Six compositional
groups were identified within the 50 thermal water samples – Group I: Na-Cl
(12 samples), Group II: Na-HCO3-Cl + Ca-SO4 (13 samples), Group
III: Na-HCO3-Cl (3 samples), Group IV: Ca-Cl (1 sample), Group V:
Ca-SO4 (4 samples), and Group VI: Ca-HCO3 (17 samples) (Table 2).
These groups likely reflect differences in sources of water, water–rock
interactions, and structural control of the local geothermal systems.
The Na-Cl and Ca-SO4 (Group I and V, respectively) type waters may have
originated via the water–rock interactions involving pockets of evaporites
in the area. Oriel and Platt (1980) have reported the presence of evaporites
(e.g., halite, gypsum, and alum) in Middle Jurassic sequences (Preuss
Redbeds) in southeastern Idaho. Recently, Welhan et al. (2014) indicated
that the high-salinity waters in some deep wildcat petroleum wells may be
related to either dissolution of salts from the Preuss evaporites or
magmatic waters from a zone as deep as 12–14 km under the BVF in the
fold–thrust belt. However, all Na-Cl (Group I) type waters considered in
this study are from the surface expressions (hot/warm springs) or from
shallow wells and may have originated via water–rock interactions involving
evaporites. This type of water is also reported from the Raft River
geothermal area (RRGA) located to the west of the present study area (Ayling
and Moore, 2013). All Ca-SO4 (Group V) type waters are from hot springs
near Bear Lake, located near the Idaho–Wyoming–Utah border. Deep sourced water
from a nearby deep wildcat petroleum well (N Eden Federal well with depth
> 2500 m) has a very high SO4 concentration; however, this
water has low Ca concentration and high Na concentration (Souder, 1985). The
Ca-SO4 (Type V) type waters that the Bear Lake hot springs issue may
have separate sources of Ca and SO4, or there may be cation exchange
reactions involving Ca and Na along the flow path.
The Ca-HCO3 (Group VI) type waters are scattered throughout the area.
These waters typically exhibit low Cl concentrations (Table 2). With some
exceptions (e.g., Black River warm spring, Corral Creek wells, Soda Geyser,
Pescadero Warm Spring), these waters also have a low SO4 concentration.
This type of water is generally regarded as a product of the interaction of
groundwater with Ca-rich rocks at shallower depth. In the adjoining eastern Snake River Plain (ESRP),
the Ca-HCO3 (Group VI) type water represents the water in the active
part of the ESRP aquifer whereas the deeper waters in ESRP area are
Na-HCO3 type (Mann, 1986; McLing et al., 2002).
Only one sample represents Ca-Cl (Group IV) type water, the Rockland W-2
well located in the westernmost part of the study area (Fig. 1). Although
this water has some similarity with the Ca-HCO3 (Group VI) and
Ca-SO4 (Group V) types of water in terms of high Ca content compared to
its Na and K concentrations, its high Cl and low Na concentrations make it
difficult to assign it as a direct product of a particular type of
water–rock interaction.
The remaining two types of waters are mixed waters – Na-HCO3-Cl (Group III) and
Na-HCO3-Cl + Ca-SO4 (Group II). The cluster
analysis did not identify a separate Na-HCO3 type water, as it is a
representative of deep waters in the adjoining ESRP. It is more likely that
these waters are Na-Cl type waters that have interacted with carbonate with
or without gypsum/anhydride.
Southeastern Idaho waters from hot springs and wells plotted on
Giggenbach diagram (Giggenbach, 1988). All partially equilibrated waters are
of Na-Cl (Group I) type waters.
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Southeastern Idaho geothermal reservoir temperatures
Giggenbach diagram
When plotted on a Giggenbach diagram (Giggenbach, 1988), the majority of the
southeastern Idaho waters in this study are in the immature zone with some
waters in the zone of partial equilibration (Fig. 3). The partially equilibrated waters in
Fig. 3 are from hot springs and wells near Preston, Idaho (Battle Creek and
Squaw hot springs), and distribution of these waters on Giggenbach diagram
indicates that these waters could have interacted with rock at a temperature
range of 260–300 ∘C. The immaturity reflects their low Na
content as well as their higher Mg content (Giggenbach, 1988). Waters with
high Mg are deemed to be unsuitable for some traditional solute
geothermometry, although there have been efforts to develop a Mg correction
(e.g., Fournier and Potter, 1979).
Temperature (∘C) estimates for southeastern Idaho thermal waters RTEst
and other geothermometers.
Springs/wellsa
Tb±σc
Quartzd
Chalcedonye
Silicaf
Na-K-Cag
Typesh
Woodruff HS
97±3
78
47
49
56
I
E. Bingham W
161±4
117
88
88
193
Squaw HS-1
179±9
151
125
123
204
Squaw HS-2
157±6
151
125
123
183
Squaw HS W-1
175±5
152
127
125
229
Squaw HS W-2
174±6
150
124
122
217
Squaw HS W-3
171±7
156
132
129
216
Battle Creek HS-1
169±5
142
116
114
205
Battle Creek HS-2
175±6
141
115
113
215
Battle Creek HS-3
170±5
142
116
114
202
Battle Creek HS-4
171±4
136
109
107
204
Wayland HS-1
175±5
125
97
97
230
Alpine WS
98±9
92
61
63
92
II
Wayland HS-2
144±7
114
85
85
84
Treasurton WS-1
111±3
105
76
77
78
Treasurton WS-2
111±9
105
76
77
113
Cleavland WS
119±7
112
83
84
106
Maple Grove HS-1
126±4
106
77
78
97
Maple Grove HS-2
123±4
128
101
100
73
Maple Grove HS-3
124±3
129
101
101
82
Maple Grove HS-4
115±7
104
74
75
54
Maple Grove HS-5
126±6
128
100
99
67
Maple Grove HS-6
122±5
115
86
87
97
Auburn HS
107±9
117
88
88
104
Johnson S
116±13
130
103
102
134
Ben Meek W-1
106±7
131
104
103
86
III
Ben Meek W-2
106±4
125
97
97
72
Ben Meek W-3
109±4
131
103
102
73
Rockland-W2
110±7
118
90
90
93
IV
Bear Lake HS-1
113±7
64
66
73
V
Bear Lake HS-2
111±7
98
68
69
94
Bear Lake HS-3
107±8
92
61
63
92
Bear Lake HS-4
121±4
86
55
57
90
Downata HS
97±3
78
47
49
49
VI
Black River WS
103±3
83
52
55
85
Pescadero WS
68±8
81
49
52
41
Henry WS
60±16
92
61
63
89
Steamboat HS
96±11
100
99
46
Soda Springs G
59±15
86
55
57
88
Lava HS-1
94±6
82
51
53
67
Lava HS-2
94±5
86
55
57
64
Portneuf R WS-1
100±6
89
59
61
92
Portneuf R WS-2
101±9
99
69
70
111
Corral Creek W-1
98±3
77
45
48
98
Corral Creek W-2
100±4
79
48
51
97
Corral Creek W-3
98±3
79
48
51
99
Corral Creek W-4
98±3
79
48
51
100
Dyer W
121±3
117
88
88
57
Anderson W
144±4
143
117
115
74
Rockland-W1
31±4
165
142
138
88
a HS: hot spring, WS: warm spring, W: well; b RTEst-estimated
temperature; cσ is standard error in each RTEst optimized
temperature; d quartz no steam loss, Fournier (1977);
e Fournier (1977); f Arnórsson et al. (1983); g Fournier and
Truesdell (1973), Mg correction applied according to Fournier and Potter II (1979);
h water types are – I: Na-Cl (12 samples), II: Na-HCO3Cl + Ca-SO4
(13 samples), III: Na-HCO3-Cl (3 samples), IV: Ca-Cl (1 sample),
V: Ca-SO4 (4 samples), and VI: Ca-HCO3 (17 samples).
Temperature estimation for Battle Creek hot spring near Preston,
Idaho. (a) The logQ/KT curves for minerals calculated using original
water chemistry with K feldspar used as a proxy for Al, (b) optimized
logQ/KT curves (FT: field temperature (43 ∘C); ET:
estimated temperature (169 ∘C), the dark horizontal bar
below ET represents the ± standard error for the estimated temperature
(±5 ∘C); cal: calcite, cha: chalcedony, dol:
disordered dolomite, mor: mordenite K, and par: paragonite).
![]()
Temperature estimated by multicomponent equilibrium
geothermometry
Estimates of reservoir temperatures for southeastern Idaho thermal waters
shown in Table 2 were made using RTEst and several conventional
geothermometers (Table 3). The RMAs that were used in RTEst consisted of
representative minerals (Mg-bearing minerals – clinochlore, illite,
saponite, disordered dolomite; Na-bearing minerals – paragonite, saponite;
K-bearing minerals – K feldspar, mordenite K, illite; Ca-bearing minerals
– calcite, disordered dolomite; and chalcedony). For thermal waters that do
not have measured Al concentration, a value determined from assumed
equilibrium with K feldspar was used in the calculations.
In MEG, the reservoir temperature is estimated by assuming a representative
RMA with which the fluid in the reservoir is believed to have equilibrated.
Next, the activities of the chemical species in solution are determined and
the saturation indices [SI =log(Q/KT)] calculated using the
laboratory measured temperature of the sample. This calculation is repeated
for a range of temperatures and the resulting SIs recalculated. The apparent
reservoir temperature is the one at which all minerals in the RMA are in
equilibrium with the reservoir fluid as indicated by near-zero logQ/KT
values on a log(Q/KT) versus temperature plot [log(Q/KT) plot]
(Reed and Spycher, 1984; Bethke, 2008). In other words, reservoir minerals
in equilibrium with the fluid at depth should yield a common equilibrium
temperature with a near-zero log(Q/KT) value for each mineral at the
same temperature; this common equilibrium temperature is assumed to be the
reservoir temperature. If all the log(Q/KT) curves do not show a
common temperature convergence, then it suggests that there exists errors in
analytical data, the selected mineral assemblage does not represent the
actual mineral assemblage in the reservoir, or the sampled water must have
been subjected to composition-altering physical and chemical processes
during its ascent from the reservoir.
Figure 4a shows typical log(Q/KT) curves of the RMA (calcite,
chalcedony, disordered dolomite, mordenite K, and paragonite) used for the
reported Battle Creek hot spring-1 water compositions. The log(Q/KT)
curves of these minerals intersect the log(Q/KT)=0 at a wide range
of temperatures, ranging from 40 ∘C (calcite) to
over 250 ∘C (paragonite), making the log(Q/KT) curves derived
from the reported water chemistry minimally useful for estimating
temperature. The range of equilibration temperature for the assumed RMA is a
reflection of physical and chemical processes that may have modified the
Battle Creek hot spring-1 water composition during its ascent to the
sampling point.
Three common composition-altering processes are degassing, mixing, and
boiling. In particular, the loss of CO2 from geothermal water due to
degassing has a direct effect on pH, and it is often indicated by
oversaturation of calcite (Palandri and Reed, 2001). Similarly, dilution of
thermal water by cooler water or enrichment of solutes by boiling is
indicated by lack of convergence of log(Q/KT) curves over a small
temperature range. In principle, these composition-altering processes can be
taken into account by simply adding them into the measured water composition
and looking for convergence of the saturation indices of the chosen mineral
assemblage, but such a graphical approach becomes cumbersome even for two
variables (e.g., temperature and CO2).
Mean and standard deviationa of estimated temperature for each
group of water.
Geothermometer
Group 1b
Group 2c
Group 3d
Group 4e
Group 5f
Group 6g
RTEst
165±22
119±11
107±1
110
113±6
98±22
Chalcedonyh
110±24
85±13
102±4
90
62±5
67±28
Na-K-Cai
196±46
91±21
77±8
93
87±10
79±21
a Standard deviation for RTEst temperatures are calculated using RTEst
temperatures of each group without incorporating standard error associated
with estimated temperature of individual sample; b Na-Cl type water
(n=12); c Na-HCO3-Cl + Ca-SO4 type water (n=13);
d Na-HCO3-Cl type water (n=3); e Ca-Cl type water (n=1),
since this water type is represented by one sample, no standard deviations
were calculated; f Ca-SO4 type water (n=4);
g Ca-HCO3 type water (n=17); h Fournier (1977);
i Fournier and Truesdell (1973), Mg correction applied according to Fournier and
Potter II (1979).
To account for possible composition-altering processes, RTEst was used to
simultaneously estimate a reservoir temperature and optimize the amount of
H2O (to account for mixing or boiling) and the fugacity of CO2 (to
account for degassing) (Palmer et al., 2014). The optimized results for Battle
Creek hot spring-1 are shown in Figure 4b. Compared to the log(Q/KT)
curves calculated using the reported water compositions (Fig. 4a), the
optimized curves (Fig. 4b) converge to log(Q/KT)=0 within a narrow
temperature range (i.e., 169±5 ∘C).
The optimized temperatures and composition parameters for the other
southeastern Idaho waters reported in Table 2 were estimated using RTEst in
the same manner. The estimated reservoir temperatures and associated
standard errors are summarized in Table 3.
Temperature estimates with traditional geothermometers
In addition to RTEst, other traditional geothermometers were also compared
(Table 3, Fig. 5). Because most of the waters from thermal springs and wells
in southeastern Idaho are geochemically immature (Fig. 3), the use of
traditional geothermometers to estimate their temperatures is unreliable.
Temperatures obtained with silica polymorphs and Na-K-Ca geothermometers can
be quite variable, compared with the RTEst temperatures. As shown in
Table 4, group-wise mean chalcedony-based reservoir temperatures are consistently
cooler than the mean RTEst-calculated reservoir temperature.
Chalcedony-based reservoir temperatures were calculated using the observed
silica concentrations. On the other hand, RTEst reservoir temperatures were
calculated with reconstructed solute concentrations (i.e., optimized for
water gain/loss). For the majority of samples, the chalcedony-based temperatures
are lower than the RTEst-estimated temperatures (Fig. 5a). In general,
wherever RTEst indicates that the sample contains an appreciable fraction of
additional water, the RTEst temperature is higher than the chalcedony-based
temperature for that sample.
Mg-corrected Na-K-Ca temperatures are relatively similar to the RTEst
temperatures; however, the trend between mean RTEst and Na-K-Ca temperature
varies with groups. In general, Na-K-Ca geothermometry predicts in cooler
temperatures in the lower temperature range and hotter temperatures in the
upper temperature range compared to the RTEst temperatures (Fig. 5b). The
main weakness of Na-K-Ca geothermometer is poor reliability in waters with a
significant amount of Mg. Even the Mg-corrected Na-K-Ca temperature
estimates have poor reliability if the Mg concentration in thermal waters
is high and is controlled by non-chlorite minerals. In southeastern Idaho,
the Mg concentration in thermal waters is likely to be controlled by
carbonate minerals (limestone/dolomite) as these waters may have interacting
with carbonate sequences in the reservoir or along the flow path. Compared
to the RTEst temperatures, Na-K-Ca temperatures are lower for all but Group
1 waters. The overprediction of temperature for Group 1 waters is likely
caused by the disproportionate (relative to Na and K) loss of Ca to
calcite precipitation as suggested by Young and Lewis (1981).
RTEst temperatures versus chalcedony (a) and
Na-K-Ca (b) temperatures for the southeastern Idaho thermal waters. Each horizontal bar
represents 1 standard error on each side in RTEst temperature estimate.
![]()
Estimated temperature versus bottom-hole temperature of wildcat petroleum
wells
As reported in Table 1, some of the wildcat petroleum wells in the fold–thrust
belt in southeastern Idaho have measured BHTs. In general, RTEst-calculated reservoir
temperatures appear positively correlated with nearby BHTs, supporting the argument
that MEG can be used to predict deep geothermal reservoir temperatures.
The North Eden Federal 21–11 well (2618 m) is located east of the Bear Lake,
near the border of Idaho, Utah, and Wyoming. This well has slightly lower BHT (84 ∘C)
than the RTEst temperature estimates (107–121 with standard error ±4 to ±8)
for nearby Bear Lake hot springs (represented by letter code BL in Fig. 1).
Similarly, the RTEst temperature estimate for Alpine Spring (letter code AL in Fig. 1)
(98±9 ∘C) is similar to the BHT of the nearest deep well, Black Mountain-1
(4158 m, 120 ∘C). However, there are some thermal features that
have different RTEst-calculated reservoir temperatures than the BHTs of some
nearby wells. For example, the BHTs measured for Federal 1–8 (167 ∘C)
and Federal 1–13 (172 ∘C) are significantly higher than the estimated
temperature (68±8 ∘C) for the closest spring (Pescadero Warm Spring
with PD letter code in Fig. 1). In this case, it is important to note that BHTs
of some nearby wells also varied slightly. For example, Bald Mountain-2 (3830 m)
has a BHT of 140 ∘C whereas the nearby Black Mountain-1 well (4158 m)
has a BHT of 120 ∘C. Such variation in temperature at depth in nearby
wells may suggest that some deep temperatures measured in wells are not equilibrated
(i.e., disturbed by drilling) or could reflect variable proximity to the
thermal water flow paths along fault or other heterogeneities.
The highest BHT was recorded for the
King 2-1 well (3927 m, 202 ∘C), but there are no RTEst
temperature estimates in the vicinity of this well. Similarly, there is no
deep measured temperature in the vicinity of Battle Creek and Squaw hot
springs near Preston. However, these hot springs discharge hot waters (up to
84 ∘C), and some of the recent shallow wells in the area
are reportedly producing water with a temperature over 100 ∘C.
New initiatives (e.g., Wood et al., 2015) would help further assess the
geothermal potential of the Preston area system.