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Gas pressures in geothermal systems
Chemical Geology, 4 9 ( 1 9 8 5 ) 3 1 9 - - 3 2 8
319
Elsevier Science P u b l i s h e r s B.V., A m s t e r d a m - - P r i n t e d in T h e N e t h e r l a n d s
GAS PRESSURES IN GEOTHERMAL SYSTEMS STE FIi.NA R N O R S S O N Science Institute, University of Iceland, 107 Reykjavik (Iceland)
(Accepted for publication July 16, 1984)
Abstract A r n 6 r s s o n , S., 1 9 8 5 . Gas pressures in g e o t h e r m a l systems. In: Y. K i t a n o ( G u e s t - E d i t o r ) , W a t e r - - R o c k Interaction. C h e m . Geol., 49: 3 1 9 - - 3 2 8 .
The partial pressures of CO2, H2S and H2 in geothermal reservoir waters are fixed by temperature-dependent minerhl equilibria. The dissolved gases have a significant effect at which depth the rising hot waters start to boil, if their temperature exceeds 250°C and, particularly, when it exceeds 300°C. The following functions describe the temperature dependence of PCo2, PH2S and PH2, respectively: log PCO 2
= - 2 . 8 1 - 5 0 1 2 . 7 T -1 - 0 . 0 0 9 1 9 T + 6 . 4 6 4 log T
logPH2s(s)
=-18.75-24738.9T
log PH2S(d)
= - 2 . 7 6 - 5 7 5 8 . 2 T -~ - 0 . 0 0 8 5 0 T + 6 . 3 5 9 log T
log PHi(s)
= - 1 0 . 1 4 - 1 7 7 6 3 . 1 T -1 - 0 . 0 4 0 8 4 T + 2 3 . 3 6 8 log T
logPH:(d )
= 9.85 + 7 2 9 0 . 3 T -~ + 0 . 0 7 2 0 2 T -
-l-0.10133T+43.1701ogT
2 2 . 6 5 1 log T
w h e r e P is partial pressure (in bars abs.); a n d T is t e m p e r a t u r e (in kelvins). T w o f u n c t i o n s are given for each o f H2S a n d H2. One set ( i n d i c a t e d b y s u b s c r i p t d) is valid for all w a t e r s < 200°C a n d w a t e r s in t h e r a n g e 2 0 0 - - 3 0 0 ° C if c h l o r i d e c o n c e n t r a t i o n s are < 5 0 0 p p m . T h e o t h e r set ( i n d i c a t e d b y s u b s c r i p t s) is valid f o r all w a t e r s > 300°C a n d w a t e r s in t h e r a n g e 2 0 0 - - 3 0 0 ° C if c o n t a i n i n g > 5 0 0 p p m C1-. I t is c o n s i d e r e d t h a t t h e t e m p e r a t u r e f u n c t i o n s are valid for w a t e r s in basaltic t o acidic volcanic r o c k s a n d also for s o m e waters, a t least, in s e d i m e n t a r y rocks. T h e f o l l o w i n g f u n c t i o n describes c r u d e l y t o t a l gas pressure plus PH:O in g e o t h e r m a l s y s t e m s : log Ptotal = 0.00543-
1. Introduction
Extensive studies of many drilled geothermal fields in the world have shown that temperatures frequently closely follow the boiling point curve for pure water with depth (see e.g., Stef~nsson and BjSrnsson, 1982). Yet geothermal waters contain dissolved gases which will, depending on their concentrations, affect the boiling point. Rising hot 0009-2541/85/$03.30
6 2 1 . 4 T -1
waters begin to boil when steam pressure plus total gas pressure become equal to the hydrostatic pressure. Giggenbach (1981) demonstrated that the partial pressure of CO: in some New Zealand geothermal systems were uniquely related to aquifer temperature and revealed how CO2 affected the boiling point. It is clear from the work of Nehring and d'Amore (1981) on the Cerro Prieto field in Mexico that hydrogen
© 1 9 8 5 Elsevier Science P u b l i s h e r s B.V.
320
and hydrogen sulphide partial pressures are fixed by temperature-dependent mineral equilibria. The present contribution is based on data from 29 hot-water wells and 20 wet-steam wells from most of the drilled geothermal fields in Iceland (Fig. 1), as well as 22 wetsteam wells in other countries. Aquifer temperatures for the wells are in the range 22-325°C. Analytical data for major components of all the Icelandic waters are tabulated in Arn6rsson et al. (1983) and Arn6rsson and Gunnlaugsson (1985). The results presented here show that the partial pressures of CO2, H2S and H~ vary regularly over a large span of aquifer temperatures. The dissolved gases have a significant effect at which depth the rising hot waters start to boil, if their temperature exceeds 250°C and, particular-
61,8"
ly, when it exceeds 300°C. Functions are presented which describe the temperature dependence of the partial pressure of CO~, H2S and H2 in geothermal reservoir waters. 2. Calculation o f gas p r e s s u r e s The basic data used to derive the gas partial pressures in the geothermal systems originate from analyses of water and steam from wetsteam well discharges and also from hot-water well discharges in the case of CO2 and H~S. For the Icelandic data Arn6rsson et al. (1983) have demonstrated that these discharges represent equilibrated waters. It is assumed that a single liquid water phase exists in the aquifer. The total discharge of wet-steam wells is taken to represent the composition of the water in this
56 •
63
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•
54
5"1
KRAFLA
NAMAFJALL 50
"1.2. /N ~ SVARTgE~I-
70
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~43I! A' ~
11
"47
48,49
HVERAGERDI 0
,
1,00kin
Fig. 1. L o c a t i o n o f s a m p l e p o i n t s in Iceland w i t h i n d e x numbers c o r r e s p o n d i n g to the first c o l u m n in Tables I a n d II. W h e n d a t a f r o m several wells in t h e s a m e field were c o n s i d e r e d t h e field n a m e is s h o w n .
321
aquifer, if boiling starts within the well, or if the measured discharge enthalpy is close to that of steam-saturated water at the aquifer temperature. In such instances the concentration, mi, d, of gas i in the aquifer was calculated from the following equation: mi, d = mi, w(1 - X )
+ mi, v X
(1)
where the subscripts d, w and v denote the aquifer water, and the water and steam samples, respectively, o f the discharge collected at a known separation pressure at the wellhead. X represents the steam fraction in the discharge, and: X = (hd, w - h w ) / L w where hd, w is the total discharge enthalpy; and hw and Lw the enthalpy of the sampled water and its latent heat of vaporisation, respectively. For all gases except CO: and H2S, mi, w is taken to be zero, as they are not analysed in the aqueous phase. Some of the wells included in this study have a "high" discharge enthalpy. For these wells it was assumed that the "excess" steam resulted from evaporation of degassed pore water by flow of heat from the rock. "Excess" steam designates steam in well discharges in excess of that formed b y adiabatic boiling o f water from a specified aquifer temperature. It is visualized that one of two processes may contribute to excess enthalpy. One is the separation of steam from water in the aquifer and preferential movement to the well and the other evaporation of pore water. When boiling is enhanced in aquifers b y pressure drop around producing wells, the water--steam mixture will cool and thermal disequilibrium b e t w e e n the rock and fluid results. Flow of heat from the rock to the fluid will increase the steam/water ratio in that fluid and it seems most likely that this will occur by evaporation of water trapped in the pores of the rock because the heat transfer will be most effective into that water. During the initial stages of boiling the pore water will b e c o m e degassed and subsequent evaporation will lead to the addi-
tion of steam free of gas to the steam--water mixture flowing towards the well. The following equation was used to derive gas concentrations, mi, d, in the reservoir water when using data from "high" enthalpy wells (in deriving this equation it was assumed that the enthalpy of the added steam is equal to that of saturated steam at the sampling pressure): mi, d = mi, w ( l - X ) + mi, vXs
(2)
where Xs is the fraction of steam in the total discharge (the difference between Xs and X represents evaporated pore water). For hot-water wells the composition o f the discharged water is taken to represent that of the deep aquifer water. It is certainly possible, even likely, that the "excess" enthalpy of some of the wet-steam well discharges is in part due to partial separation of water and steam in the aquifer, due to the relative permeability effect. If phase separation was the only cause of the high discharge enthalpy and it occurred at a pressure close to the sampling pressure, then eq. 1 describes correctly the relation between the analytical data and the deep unboiled water gas concentrations. The difference in the calculated values for mi, d by eqs. 1 and 2 depends on the discharge enthalpy value in relation to the enthalpy of steam-saturated water at the aquif6r temperature. The arrows in Fig. 2 give an impression of the magnitude of the difference for selected wells. Inspection shows that the overall pattern of the gas partial pressure--aquifer temperature relationship, discussed in Section 3, is not changed by selection of either o f the t w o models described above to explain "excess" enthalpy well discharges. Nehring and d'Amore ( 1 9 8 1 ) d o not report discharge enthalpy for the Cerro Prieto wells used for the present study, but take it to be equal to that o f steam-saturated water at the Na--K--Ca geothermometry temperature. This corresponds to the use of eq. 1 in calculating mi, d for the various gases. Gas pressures were obtained from the cal-
322 culated gas concentrations in the deep aquifer water using the values of the respective Henry's law coefficients reported by Naumov et al. (1971). Activity coefficients for gases were taken to be equal to unity. The program of ArnSrsson et al. (1982) was used to compute the deep aquifer composition and the gas pressures, taking into account chemical speciation in the case of carbonate and sulphide. However, in the case of the New Zealand and Cerro Prieto data (see Table I) information was inadequate for speciation computations and analysed total carbonate and total sulphide were taken to be equal to CO: and H:S. Inspection of other data shows that CO: exceeds 80% o f total carbonate for waters of > 240°C and is 98--99% for waters of ~ 300°C. Comparable figures for sulphide are more than 70% and 97--99%, respectively. It appears, therefore, that the approximation taking total carbonate and total sulphide to represent CO: and H:S respectively for the New Zealand and Cerro Prieto data does not yield significantly high results. Gas pressures have also been evaluated after adiabatic boiling o f selected waters to various preset temperatures, again using the c o m p u t e r program o f Arn6rsson et al. (1982). 3. Gas pressures in reservoir waters
The aquifer partial pressures of CO2, H2S, CH4 and H2 in m a n y geothermal systems are summarized in Tables I and II and in Fig. 2. Pco~ displays a particularly good relationship with temperature, both for wetsteam and hot-water well data, to temperatures as low as 22°C. The scatter for PH~S and PH~ is much greater. Inspection o f the data points in Fig. 2 shows a tendency for PH~s and PH~ to fall on one o f two curves. One curve corresponds with all waters of ~ 200°C and waters in the range 200--300°C when containing ~ 500 ppm C1-. The other curve corresponds to all waters of > ~ 300°C and waters in the range 200--300 ° when contain-
ing > 500 ppm C1-. There is practically no overlap between data points belonging to either of the two curves so defined. According to ArnSrsson and Gunnlaugsson {1985) the two different curves for PH~S and PHi, are due to buffering of PH~S and PHi, respectively, by different mineral assemblages. Above 230°C and for "dilute" water the assemblage includes pyrite, pyrrhotite, epidote and prehnite, but pyrrhotite, epidote, prehnite and magnetite or chlorite are involved in the case o f the "saline" waters. PCH4does not show any clear dependence on aquifer temperature (Fig. 2) which would be expected at equilibrium. Examination shows that the "dilute" geothermal waters and waters of < 200°C depart considerably from equilibrium for the Fischer--Tropsch reaction: CO2 + 4H2 ~- CI-h + 2H20
(3)
Higher methane and/or lower hydrogen gas pressure would change the system towards equilibrium. By contrast, some of the "saline" waters, particularly those from New Zealand, closely approach equilibrium, a result which conforms with that o f Giggenbach (1980). Other saline waters depart considerably from equilibrium at the aquifer temperature: the Cerro Prieto waters contain less methane than would be required for equilibrium whereas the situation is the reverse for the Svartsengi and Reykjanes geothermal seawaters in Iceland. With few exceptions (some of the hottest waters) PN~ is below 2--3 bar (Fig. 2), which is somewhat higher or equal to the level anticipated if the sole source o f the nitrogen was atmospheric and dissolved in the meteoric source water (~ 0.5 umol N2 per kg H20). Above ~ 250°C Pco: is significant relative to PH~O and will thus have a significant influence at which total pressure the geothermal waters will begin to boil. The influence of Pco~ rises abruptly with temperature. At 300 C it is ~ 10 bar (abs.) and is probably as high as 30 bar (abs.) at 350°C. PH2 also significantly affects the boiling point for "dilute"
323
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i o'
o~o~ e
•'@
o
@
,,#o ~ '
o
lc;o
'
26o
'
• =- '
300
160
Aquifer temp, °C
ffi .1 ,ll
E
•
oo
• li•
0
•
o~O
o 0
o.
•
o
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D
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-3
260
O
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• o •
HOT-WATER WELLS:
•
•
35o
o
21~0
•
0
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ICELAND
i
260
3(:)0
m WAIRAKEI, KAWERAU, BROADLANDS; NEW ZEALAND
• ONUMA; •
%
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....
-2
Aquifer temp. °C
WET-STEAM WELLS: ICELAND
o o
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a
3(~0
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Aquifer temp. °C
08
oan~O •
_o
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¢'1
,,o
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260
4~1 .1
~-1
•
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Aquifer temp. °C
'
3~o
Aquifer temp. °C JAPAN
OLKARIA; KENYA
V TONGONAN, PALINPINON; PHILIPPINES
o CERRO PRIETO; MEXICO
Fig. 2. Partial pressure for CO2, H2S, H 2, CH 4 and N2 in geothermal waters. The curve on the bottom righthand diagram representsPN~ corresponding to a concentration o f 0.5 u m o l , the quantity dissolved in meteoric water• Open symbols refer to all @aters at and above 300°C and waters in the range 200--300°C, if containing more than 500 p p m C1-. Filled symbols designate all waters below 200°C and waters in the range 200--300°C if containing less than 500 ppm C1-. The arrows in the top two diagrams o f the figure indicate the shift in the respective gas partial pressure if eq. 1, rather than eq. 2 was used to derive gas concentration in the deep reservoir water (see text)•
waters, i.e. waters containing less than 500 ppm C1- at temperatures above ~ 250°C. For "saline" waters PH2 is alSO significant above 300°C. Other gases do not exert a significant partial pressure relative to steam pressure, except N2, particularly in the case of waters from Krafla, northern Iceland. Here PN~ is as high as 40 bar (abs.). Although some o f the hottest waters contain the highest nitrogen, its activity is not predictable in terms of temperature. PH~S is below 1 bar for all waters, as well as PcH4 in the case of the "dilute" waters. For waters from Cerro Prieto (Mexico), and Kawerau and Broad-
lands (New Zealand), the methane partial pressure approaches and somewhat exceeds 1 bar (abs.). Fig. 3 shows the total pressure at which individual waters listed in Table I will begin to boil and h o w this boiling point relates to that of pure water. The following equation describes the data points in Fig. 3: l o g Ptotal = 0 . 0 0 5 4 3 T -
6 2 1 . 4 T -1
(4)
where P is in bars (abs.); and T in kelvins. In calculating the total pressure, which was done with the use of the program of ArnSrsson et al. (1982), PH~O was taken to be
324
TABLE I Gas partial pressures in geothermal reservoir water -- Data from wet-steam wells Location
1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20
Reykjanes 8, Ice. Svartsengi 4, Ice. Svartsengi 5, Ice. Svartsengi 6, Ice. Hveragerdi 2, Ice. Hveragerdi 4, Ice. Hveragerdi 6, Ice. Hveragerdi 7, Ice. Nesjavellir 5, Ice. Nesjavellir 6, Ice. ReykjabS1 1, Ice. Rey kholt 1, Ice. N~imafjall 8, Ice. N~mafjall 11, Ice. Krafla 6, Ice. Krafla 8, Ice. Krafla 9, Ice. Krafla 13, Ice. Krafla 14, Ice. Krafla 16, Ice.
21
Onuma O-6R, Japan
22 23 24
Olkaria 6, Kenya Olkaria 10, Kenya Olkaria 12, Kenya
25 26 27 28 29
Palinpinon 1, Phil. Palinpinon 7, Phil. Tongonan 202, Phil. Tongonan 404, Phil. Tongonan 407, Phil.
30 31 32 33 34 35 36 37 38
Wairakei 72, N.Z. Wairakei 81, N.Z. Kawerau 7, N.Z. Kawerau 8, N.Z. Kawerau 17, N.Z. Kawerau 19, N,Z. Broadlands 11, N.Z. Broadlands 23, N.Z. Broadlands 25, N.Z.
39 40 41 42
Cerro Cerro Cerro Cerro
79-009 79-008 79-139 79-138 79-136 79-032 81-016 81-017 72-146 83-013 79-030 79-028 79-052 81-022 81-020 81-029 79-048 81-033 81-035 81-036
CO 2
Prieto Prieto Prieto Prieto
Aquifer discharge depth .1 of:
log(partial pressure in bars abs.)
No. (for Sample Iceland No. see Fig. i)
5, Mex. 8, Mex. 35, Mex. 42, Mex.
H2S
enthalpy (J g - ' )
inflow (m)
248 *2 240 240 240 182 181 215 225 271 290 152 133 246 320 300 215 240 325 295 300
1,151 1,038 *3 1,038 *3 1,038 *3 772 *3 768 *3 921 *3 967 *3 1,190 *~ 2,062 641 *3 559 *3 1,093 2,355 1,940 921 *3 1,038 *9 1,877 2,663 1,607
1,010--1,740 1,090--1,565 835,860
CH~
0.100 - 1 . 7 5 4 0.017 - 2 . 2 4 0 -0.087 -2.319 0.079 - 1 . 9 0 7 -0.796 -1.924 -0.672 -1.955 -0.324 -1.670 -0.166 -1.504 0.841 - 0 . 5 2 9 0.610 - 0 . 0 9 5 -1.290 -3.012 - 1 . 9 4 7 -3.801 0.199 - 0 . 5 2 4 0.450 - 0 . 2 1 8 1.117 - 1 . 1 5 6 -0.111 -1.363 0.164 - 1 . 2 8 2 0.905 - 0 . 8 5 4 1.668 1.054 1.037 - 0 . 5 5 9
-1.996 -2.495 -2.237 -2.521 -1.422 1.337 -1.117 -1.169 0.502 0.867 -1 .857 -1.936 0.871 1.061 0.000 -1 .583 -0.597 0.164 0.873 0.336
-2.129 -1.446 -2.331 -1.306 -2.573 -0.695 -2.219 -0.412 -2.291 -0.815 -2.191 -0.703 -2.141 0.173 -1.943 0.107 -0.374 0.170 -1.479 -0.208 -1.900 0.491 -2.007 0.520 -0.072 0.143 -1.424 1.326 -0.762 1.104 -1.241 - 0 . 5 4 5 -1.363 -0.638 -1.382 0.464 0.590 1.576 -0.955 0.158
-0.111
-1.754
-1.996
-2.307
0.185
210
0.764 0.491 0.439
-0.362 -0.896 -0.870
0.365 0.190 0.336
-0.175 -0.764 -0.818
-0.251 -0.304 -0.418
242 261 254
2,281 2,355 2,240
0.734 1.260 0.491 1.407 1.288
-1.457 -0.845 -1.426 -0.836 -0.939
277 318 312 313 315
1,147 1,326 1,395 2,065 1,770
248 *2 244 *2 265 *2 250 *~ 261 *2 269 *2 284 *2 286 *2 298 *5
1,295 1,350 1,030 1,030 1,030 1,210 1,080 1,265 1,390
289 *2 290 *2 293 *2 281 *2
-----
0.203 - 1 . 7 1 9 -0.558 -1.828 0.699 - 1 . 2 0 6 0.878 - 1 . 1 7 4 0.650 - 1 . 2 8 4 0.947 - 0 . 9 4 8 0.961 - 1 . 4 7 7 1.404 - 1 . 0 1 4 1.353 - 1 . 1 8 7 0.694 1.071 0.941 0.771
-0.844 -0.998 -1.119 -0.867
-1 .731 -2.829 -1.759 -1.202 -1.466 -0.993 -1.454 -0.837 -0.873 -0.016 -0.352 -0 .171 -0.246
N2
temperature (°C)
H2
-0.937 -0.840 -2.838 -1.290 -0.889 -0.678 0.074 - 0 . 3 4 0 -0.059 -0.497 0.170 - 0 . 2 3 9 -0.029 -0.431 0.370 - 0 . 0 9 2 0.353 - 0 . 0 1 0 0.243 0.008 0.011 0,257
-0.759 -1.055 -0.882 -0.764
4OO 630--670 64O 405 980, 1,220 1,096 773 745 850, 1,000 1,430 1,100, 1,520 1,200 1,226 1,650 1,030, 2,030 1,050 1,078, 1,274 8O0 728, 1,025 .4 750, 855--880 2,650 2,600 1,100--1,250 1,500--1,600 1,325--1,405
• 1 More detailed data for 79 and 81 samples are given by Arn6rsson et al. (1983) and Arn6rsson and Gunnlaugsson (1985), respectively. • 2 Derived using the Na--K--Ca geothermometer. • 3 Calculated from the aquifer temperature. Boiling starts within the well. It is therefore safe to assume the discharge enthalpy to be equal to that of steam-saturated water at the respective aquifer temperature. • 4 Dominant aquifer. The data from Onuma, Japan; Olkaria, Kenya; Palinpinon and Tongonan, Philippines; the New Zealand fields; and Cerro Prieto, Mexico, were derived from the following sources: H. Sakai (pers. commun., 1982); Muna (1982); Jordan (1982); Baltasar (1980); Giggenbach (1980); a n d Nehring and d'Amore (1981), respectively.
325
TABLE II Carbon dioxide and hydrogen sulphide partial pressures in hot-water well discharges in Iceland No. (see Fig. 1 )
Sample Location No.
log PCO 2
log PH~S
Aquifer temperature
Fluid inflow (m)
(°c) 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61 62 63 64 65 66 67 68 69 70 71
79-034 79-033 79-010 79-001 79-012 79-015 81-028 79-050 79-046 81-013 81-014 79-044 81-011 81-015 79-042 79-041 79-040 81-006 79-039 81-005 79-038 81-003 79-057 79-056 79-022 79-020 79-019 79-016 79-017
Bakki 1 Oxnalaekur 1 .~rbaer 1 KlausturhSlar 1 HfisatSttir 4 Laugaland 3 Laugaland 4 Urridavatn 3 Laugar 2 Gr:~tubakki 1 Svalbardseyri 2 Laugaland 7 Gler~rgil 7 TjSrn 1 Dalv~ 10 01afsfjSrdur 3 Siglufj6rdur 7 Reykir 1 Saud~rkrSkur 1 Saud~rkrbkur 12 Varmahlfd 1 Laugarbakkar 3 ReykhSlar 2 Saelingsdalur 10 Baer 3 Seltjarnarnes 4 Reykjav~ 11 Reykir 17 Reykir 5
-2.503 -0.943 -3.397 -0.409 -3.370 -4.260 -3.754 -4.770 -5.180 -5.714 -4.893 -3.697 -4.747 -5.476 -4.804 -4.616 -4.429 -4.810 -4.427 -4.569 -3.257 -3.003 -2.860 -4.268 -2.939 -2.936 -2.921 -3.851 -3.509
-3.883 -2.947 -5.556 -2.921 -6.390 -6.857 -5.896 -7.069 -7.420 -7.914 -7.197 -6.096 -7.013 -8.021 -7.038 -7.145 -6.842 -6.951 -5.943 -6.046 -4.656 -5.037 -4.987 -5.870 -5.058 -4.503 -4.863 -6.253 -5.449
134 157 86 160 70 61 96 59 64 22 54 93 55 34 64 67 67 58 68 70 92 102 100 61 109 114 129 76 81
475--871.1 936 451 80, 120 .1 170 105 *2 750 170 250 .2, 522 *2 65 970, 1,330, 1,490 1,120, 1,440, 1,485 330, 450, 740 100 818 1,128 550 590 ? 260--471.1 83, 183 -23--285 113, 120 245, 275, 320 195--1,945 .1 927 1,750 *2 270--1,045 .1
• 1 Several aquifers were encountered in the depth range indicated. The temperature between the hottest and coldest aquifer is always less than 10°C. .2 Dominating aquifer(s). Analyses of waters collected in 1979 and 1981 are given by Arnbrsson et al. (1983) and Arnbrsson and Gunnlaugsson (1985), respectively. They also give more detailed information on physical drillhole data.
equal to that of steam-saturated water at the respective aquifer temperature. This is a satisfactory approximation when considering the scatter of data points for gas partial pressures and because PH20 is not so sensitive to total pressure for the range in question and in the presence of liquid water. Eq. 4 describes approximately the temperature--pressure relationship in boiling geothermal systems which have closely ap-
proached solute--mineral equilibrium. The equation is only approximate, since PN2 varies irregularly with temperature and because PH2S and PH2 for both "saline" and "dilute" waters are combined into one equation. Eqs. 5--9 describe the temperature dependence of Pco2, PH2Sand PH2: log Pco2
= -2.81 - 5012.7T -~ - 0.00919T + 6.464 log T
(5)
326
the range 200--300°C if they contain more than 500 p p m C1-. Eqs. 7 and 9 (with subscripts d) are valid for all waters below 200°C and for waters in the range 200--300°C if they contain ~ 500 ppm C1-. Eq. 5 describing the temperature dependence of Pco~ differs somewhat from that given b y Giggenbach (1981), b u t he used data largely in the range 200--300°C. Eq. and his equation yield 0.35 and 0.05 bar (abs.) at 200°C, practically the same at 250°C, but the difference at 300°C is 27 bar (abs.). Visual inspection of the data points for Pco~ in Fig. 2 above 200°C does not reflect an unambiguous relation with aquifer temperature, which has recently been a source of disagreement (Giggenbach, 1982; Grant, 1982). However, b y considering waters over a larger range of temperatures, as is the case for the present study (see Fig. 2), such dependence is evident.
50 .Q
XI
u~ u) L. Q.
~o
!
150
KRAFLA WELL 14
4. Boiling
200 300 Temperature °C
Fig. 3. Relation between total pressure (steam and gases) and aquifer temperature in geothermal waters. The broken line inducates temperature---pressure relationship for water from well 14 in Krafla during adiabatic boiling and equilibrium degassing.
log Pg~s(s) = - 1 8 . 7 5 - 24738.9T -1 - 0 . 1 0 1 3 3 T + 43.170 log T
(6)
log PH~S¢d) = - - 2 . 7 6 - 5 7 5 8 . 2 T -1 - 0.00850T + 6.359 log T logPH~(s )
=-10.14-
1ogPH~(d)
(7)
1 7 7 6 3 . 1 T -~
0 . 0 4 0 8 4 T + 23.368 log T
(8)
= 9.85 + 7290.3T -1 + 0 . 0 7 2 0 2 T -
22.651 log T
(9)
where P is again in bars (abs.); and T in kelvins. Eqs. 6 and 8 {with subscripts s) are valid for all waters of > 300°C and for waters in
point
curves
Boiling of initially equilibrated geothermal water in a closed system will eventually lead to almost quantitative transfer of the dissolved gases into the steam and, as a result, the gas partial pressures will be drastically reduced. This implies that a large pressure drop will accompany a small temperature drop during the early stages of boiling, if gas partial pressures are significant relative to the vapour pressure. Under closed-system conditions each geothermal water will possess a unique boiling point curve dictated b y initial temperature and its gas composition. One such curve is shown in Fig. 3 for well 14 in the Krafla geothermal field and in Fig. 4 for other selected wells in terms of drop in total gas pressure for adiabatic boiling. The program of Arn6rsson et al. (1982) was used to calculate these curves, simultaneously taking into account the distribution of eight gases (CO2, H2S, H2, CH4, N2, 02, Ar and NH3) between the water and steam phases. The initial slope of the boiling point curves will depend on which gas dominates
327
the pressure, being steepest when the least soluble gases, H2 and N2, contribute most to the total gas pressure. As the boiled water becomes more and more depleted in total gas with continued steam formation all boiling point curves approach more and more closely the boiling point curve for pure water. Degassing accompanying boiling in geothermal systems m a y often be incomplete. However, this has little effect on the shape of the boiling point curves, even if the degas1 sing is as little as ~ of m a x i m u m (that is equilibrium degassing). The reason is that all the gases are strongly partitioned into the steam phase. Salinity, which will reduce gas solubility, also has little effect on the shape of the boiling point curve. i
Krafla well 16
101 .,0
Hveragerdi well 7
,//
/
..D
ffl In
Je It:e""
~. 10-1 O)
10-3
;/ I
I
200
I
I
300 Temperature °C
Fig. 4. Reduction in gas pressure for selected geothermal waters during adiabatic boiling and equilibrium degassing.
5. Discussion
Most of the data considered in the present study provide reliable data on gas partial pressures, namely data from hot-water and
wet-steam wells where the boiling starts within the well. The data from "high-" enthalpy wells might be regarded as suspect and the gas content of the well discharge may not correspond with that dissolved in the water at depth, as indicated b y eq. 2. When comparing these two types of data it is seen that they give a coherent picture of the relation between gas partial pressure and aquifer temperatures (Fig. 2). From this it is concluded that the model adopted and described in a previous section is satisfactory, as well as the temperature dependence of gas partial pressures expressed by eqs. 5--9. It is evident that the same mineral assemblages will not buffer the gas partial pressures over the temperature range of 22--325°C considered in this study. Yet a smooth temperature dependence is observed for Pco~, the reason for this most likely being that the free energy difference between the different assemblages is small. Such is often known to be the case for metamorphic mineral assemblages. The same does not, however, hold for PH2S and PHi, as the different mineral buffers fix the respective gas pressures at significantly different levels at any particular temperature. It has not been established which minerals constitute these buffers at low temperatures, but at elevated temperatures (above ~ 200--250°C) Giggenbach {1981) and Arn6rsson and Gunnlaugsson (1985) have evaluated which assemblages are most likely to be involved. It is often observed that temperature profiles in geothermal wells deviate from the boiling point curve of pure water to lower temperatures for the measured total pressure. This deviation may in some instances, particularly at temperatures above 300°C, be due to the effect of gas pressure on the boiling point. In other wells slow heating-up may eventually bring temperatures to the boiling point for pure water. Here the reason m a y be that boiling and degassing occurs in the closed well, as evidenced b y accumulation of gas under the wellhead. The flow m a y either be provoked by condensation o f steam at
328 high levels in t h e well or b y f l o w o f fluid into a s h a l l o w e r aquifer. With t h e aid o f eqs. 5 - - 9 , analyses o f t h e c o m p o s i t i o n o f t h e gas u n d e r the wellhead and direct measurement of total gas p r e s s u r e ( t o t a l p r e s s u r e m i n u s s t e a m pressure as i n d i c a t e d b y t e m p e r a t u r e ) it m a y be possible t o e v a l u a t e t h e t e m p e r a t u r e o f t h e d o m i n a t i n g a q u i f e r in t h e well.
0° to 370°C. Geochim. Cosmochim. Acta, 46: 1513--1532. Arn6rsson, S., Gunnlaugsson, E. and Svavarsson, H., 1983. The chemistry of geothermal waters in Iceland, II. Mineral solubilities a n d independent variables controlling water composition. Geochim. Cosmochim. Acta, 47: 1513--1532. Baltasar, A.J., 1980. Interpretation of the water and gas chemistry from three geothermal areas in the Philippines Manito in Albany, Biliran Island and Tongonan in Leyte. U.N. Univ. Geotherm. Train. Prog., Reykjav~, Rep. 1980-3, 55 pp. Giggenbach, W.F., 1980. Geothermal gas equilibria. Geochim. Cosmochim. Acta, 44: 2021--2032. Giggenbach, W.F., 1981. Geothermal mineral equilibria. Geochim. Cosmochim. Acta, 45: 393--410. Giggenbach, W.F., 1982. Reply to a Comment by M.A. Grant. Geothermal mineral equilibria. Geochim. Cosmochim. Acta, 46: 2681--2684. Grant, M.A., 1982. On the lack of a unique relation between CO s partial pressure and temperature in geothermal systems. Comment on Geothermal mineral equilibria, by W.F. Giggenbach. Geochim. Cosmochim. Acta, 46: 2677--2680. Jordan, O.T., 1982. Implication of solution--mineral equilibria on the exploitation of the S-Negros geothermal field, Philippines. U.N. Univ. Geotherm. Train. Prog., Reykjav~, Rep. 1982-7, 67 pp. Muna, Z.W., 1982. Chemistry of well discharges in the Olkaria geothermal field, Kenya. U.N. Univ. Geotherm. Train. Prog., Reykjav~, Rep. 1982-8, 38 pp. Naumov, G.B., Ryzhenko, B.N. and Khodakovsky, I.L., 1971. Handbook of thermodynamic data. U.S. Geol. Surv., Rep. WRD-74-001 (translated of a Russian report). Nehring, N.L. and d'Amore, F., 1981. Gas chemistry and thermometry of the Cerro Prieto geothermal field. Proc. 3rd Syrup. on the Cerro Prieto Geothermal Field, Baja California, March 1981. Steffinsson, V. and Bj6rnsson, S., 1982. Physical aspects of hydrothermal systems. In: Continental and Oceanic Rifts. Geodyn. Ser., 8: 123--145. -
Acknowledgements T h i s article was originally p r e s e n t e d at the Fourth International Symposium on W a t e r - - R o c k I n t e r a c t i o n . I t is n o w p r e s e n t e d in an e x t e n d e d f o r m b y i n c o r p o r a t i n g d a t a f r o m g e o t h e r m a l fields, n o t o n l y f r o m Iceland, b u t f r o m o t h e r c o u n t r i e s as well. Dr. Einar Gunnlaugsson from the Reykjav~ D i s t r i c t H e a t i n g Service is specially t h a n k e d f o r his assistance w i t h c o m p u t e r w o r k and Dr. V a l g a r d u r Steffinsson o f t h e N a t i o n a l E n e r g y A u t h o r i t y , Reykjaw~k, f o r p r o v i d i n g s o m e o f t h e p h y s i c a l drillhole data. P r o f e s s o r H i t o s h i Sakai is also t h a n k e d f o r p r o v i d i n g c h e m i c a l a n d p h y s i c a l drillhole d a t a f o r t h e O n u m a field, J a p a n .
References Arn6rsson, S. and Gunnlaugsson, E., 1985. New gas geothermometers for geothermal exploration -calibration and application. Geochim. Cosmochim. Acta, 49 (in press). Arn6rsson, S., Sigurdsson, S. and Svavarsson, H., 1982. The chemistry of geothermal waters in Iceland, I. Calculation of aqueous speciation from
-
319
Elsevier Science P u b l i s h e r s B.V., A m s t e r d a m - - P r i n t e d in T h e N e t h e r l a n d s
GAS PRESSURES IN GEOTHERMAL SYSTEMS STE FIi.NA R N O R S S O N Science Institute, University of Iceland, 107 Reykjavik (Iceland)
(Accepted for publication July 16, 1984)
Abstract A r n 6 r s s o n , S., 1 9 8 5 . Gas pressures in g e o t h e r m a l systems. In: Y. K i t a n o ( G u e s t - E d i t o r ) , W a t e r - - R o c k Interaction. C h e m . Geol., 49: 3 1 9 - - 3 2 8 .
The partial pressures of CO2, H2S and H2 in geothermal reservoir waters are fixed by temperature-dependent minerhl equilibria. The dissolved gases have a significant effect at which depth the rising hot waters start to boil, if their temperature exceeds 250°C and, particularly, when it exceeds 300°C. The following functions describe the temperature dependence of PCo2, PH2S and PH2, respectively: log PCO 2
= - 2 . 8 1 - 5 0 1 2 . 7 T -1 - 0 . 0 0 9 1 9 T + 6 . 4 6 4 log T
logPH2s(s)
=-18.75-24738.9T
log PH2S(d)
= - 2 . 7 6 - 5 7 5 8 . 2 T -~ - 0 . 0 0 8 5 0 T + 6 . 3 5 9 log T
log PHi(s)
= - 1 0 . 1 4 - 1 7 7 6 3 . 1 T -1 - 0 . 0 4 0 8 4 T + 2 3 . 3 6 8 log T
logPH:(d )
= 9.85 + 7 2 9 0 . 3 T -~ + 0 . 0 7 2 0 2 T -
-l-0.10133T+43.1701ogT
2 2 . 6 5 1 log T
w h e r e P is partial pressure (in bars abs.); a n d T is t e m p e r a t u r e (in kelvins). T w o f u n c t i o n s are given for each o f H2S a n d H2. One set ( i n d i c a t e d b y s u b s c r i p t d) is valid for all w a t e r s < 200°C a n d w a t e r s in t h e r a n g e 2 0 0 - - 3 0 0 ° C if c h l o r i d e c o n c e n t r a t i o n s are < 5 0 0 p p m . T h e o t h e r set ( i n d i c a t e d b y s u b s c r i p t s) is valid f o r all w a t e r s > 300°C a n d w a t e r s in t h e r a n g e 2 0 0 - - 3 0 0 ° C if c o n t a i n i n g > 5 0 0 p p m C1-. I t is c o n s i d e r e d t h a t t h e t e m p e r a t u r e f u n c t i o n s are valid for w a t e r s in basaltic t o acidic volcanic r o c k s a n d also for s o m e waters, a t least, in s e d i m e n t a r y rocks. T h e f o l l o w i n g f u n c t i o n describes c r u d e l y t o t a l gas pressure plus PH:O in g e o t h e r m a l s y s t e m s : log Ptotal = 0.00543-
1. Introduction
Extensive studies of many drilled geothermal fields in the world have shown that temperatures frequently closely follow the boiling point curve for pure water with depth (see e.g., Stef~nsson and BjSrnsson, 1982). Yet geothermal waters contain dissolved gases which will, depending on their concentrations, affect the boiling point. Rising hot 0009-2541/85/$03.30
6 2 1 . 4 T -1
waters begin to boil when steam pressure plus total gas pressure become equal to the hydrostatic pressure. Giggenbach (1981) demonstrated that the partial pressure of CO: in some New Zealand geothermal systems were uniquely related to aquifer temperature and revealed how CO2 affected the boiling point. It is clear from the work of Nehring and d'Amore (1981) on the Cerro Prieto field in Mexico that hydrogen
© 1 9 8 5 Elsevier Science P u b l i s h e r s B.V.
320
and hydrogen sulphide partial pressures are fixed by temperature-dependent mineral equilibria. The present contribution is based on data from 29 hot-water wells and 20 wet-steam wells from most of the drilled geothermal fields in Iceland (Fig. 1), as well as 22 wetsteam wells in other countries. Aquifer temperatures for the wells are in the range 22-325°C. Analytical data for major components of all the Icelandic waters are tabulated in Arn6rsson et al. (1983) and Arn6rsson and Gunnlaugsson (1985). The results presented here show that the partial pressures of CO2, H2S and H~ vary regularly over a large span of aquifer temperatures. The dissolved gases have a significant effect at which depth the rising hot waters start to boil, if their temperature exceeds 250°C and, particular-
61,8"
ly, when it exceeds 300°C. Functions are presented which describe the temperature dependence of the partial pressure of CO~, H2S and H2 in geothermal reservoir waters. 2. Calculation o f gas p r e s s u r e s The basic data used to derive the gas partial pressures in the geothermal systems originate from analyses of water and steam from wetsteam well discharges and also from hot-water well discharges in the case of CO2 and H~S. For the Icelandic data Arn6rsson et al. (1983) have demonstrated that these discharges represent equilibrated waters. It is assumed that a single liquid water phase exists in the aquifer. The total discharge of wet-steam wells is taken to represent the composition of the water in this
56 •
63
_~53
5s-,
•
54
5"1
KRAFLA
NAMAFJALL 50
"1.2. /N ~ SVARTgE~I-
70
•-_ ~,,,
~43I! A' ~
11
"47
48,49
HVERAGERDI 0
,
1,00kin
Fig. 1. L o c a t i o n o f s a m p l e p o i n t s in Iceland w i t h i n d e x numbers c o r r e s p o n d i n g to the first c o l u m n in Tables I a n d II. W h e n d a t a f r o m several wells in t h e s a m e field were c o n s i d e r e d t h e field n a m e is s h o w n .
321
aquifer, if boiling starts within the well, or if the measured discharge enthalpy is close to that of steam-saturated water at the aquifer temperature. In such instances the concentration, mi, d, of gas i in the aquifer was calculated from the following equation: mi, d = mi, w(1 - X )
+ mi, v X
(1)
where the subscripts d, w and v denote the aquifer water, and the water and steam samples, respectively, o f the discharge collected at a known separation pressure at the wellhead. X represents the steam fraction in the discharge, and: X = (hd, w - h w ) / L w where hd, w is the total discharge enthalpy; and hw and Lw the enthalpy of the sampled water and its latent heat of vaporisation, respectively. For all gases except CO: and H2S, mi, w is taken to be zero, as they are not analysed in the aqueous phase. Some of the wells included in this study have a "high" discharge enthalpy. For these wells it was assumed that the "excess" steam resulted from evaporation of degassed pore water by flow of heat from the rock. "Excess" steam designates steam in well discharges in excess of that formed b y adiabatic boiling o f water from a specified aquifer temperature. It is visualized that one of two processes may contribute to excess enthalpy. One is the separation of steam from water in the aquifer and preferential movement to the well and the other evaporation of pore water. When boiling is enhanced in aquifers b y pressure drop around producing wells, the water--steam mixture will cool and thermal disequilibrium b e t w e e n the rock and fluid results. Flow of heat from the rock to the fluid will increase the steam/water ratio in that fluid and it seems most likely that this will occur by evaporation of water trapped in the pores of the rock because the heat transfer will be most effective into that water. During the initial stages of boiling the pore water will b e c o m e degassed and subsequent evaporation will lead to the addi-
tion of steam free of gas to the steam--water mixture flowing towards the well. The following equation was used to derive gas concentrations, mi, d, in the reservoir water when using data from "high" enthalpy wells (in deriving this equation it was assumed that the enthalpy of the added steam is equal to that of saturated steam at the sampling pressure): mi, d = mi, w ( l - X ) + mi, vXs
(2)
where Xs is the fraction of steam in the total discharge (the difference between Xs and X represents evaporated pore water). For hot-water wells the composition o f the discharged water is taken to represent that of the deep aquifer water. It is certainly possible, even likely, that the "excess" enthalpy of some of the wet-steam well discharges is in part due to partial separation of water and steam in the aquifer, due to the relative permeability effect. If phase separation was the only cause of the high discharge enthalpy and it occurred at a pressure close to the sampling pressure, then eq. 1 describes correctly the relation between the analytical data and the deep unboiled water gas concentrations. The difference in the calculated values for mi, d by eqs. 1 and 2 depends on the discharge enthalpy value in relation to the enthalpy of steam-saturated water at the aquif6r temperature. The arrows in Fig. 2 give an impression of the magnitude of the difference for selected wells. Inspection shows that the overall pattern of the gas partial pressure--aquifer temperature relationship, discussed in Section 3, is not changed by selection of either o f the t w o models described above to explain "excess" enthalpy well discharges. Nehring and d'Amore ( 1 9 8 1 ) d o not report discharge enthalpy for the Cerro Prieto wells used for the present study, but take it to be equal to that o f steam-saturated water at the Na--K--Ca geothermometry temperature. This corresponds to the use of eq. 1 in calculating mi, d for the various gases. Gas pressures were obtained from the cal-
322 culated gas concentrations in the deep aquifer water using the values of the respective Henry's law coefficients reported by Naumov et al. (1971). Activity coefficients for gases were taken to be equal to unity. The program of ArnSrsson et al. (1982) was used to compute the deep aquifer composition and the gas pressures, taking into account chemical speciation in the case of carbonate and sulphide. However, in the case of the New Zealand and Cerro Prieto data (see Table I) information was inadequate for speciation computations and analysed total carbonate and total sulphide were taken to be equal to CO: and H:S. Inspection of other data shows that CO: exceeds 80% o f total carbonate for waters of > 240°C and is 98--99% for waters of ~ 300°C. Comparable figures for sulphide are more than 70% and 97--99%, respectively. It appears, therefore, that the approximation taking total carbonate and total sulphide to represent CO: and H:S respectively for the New Zealand and Cerro Prieto data does not yield significantly high results. Gas pressures have also been evaluated after adiabatic boiling o f selected waters to various preset temperatures, again using the c o m p u t e r program o f Arn6rsson et al. (1982). 3. Gas pressures in reservoir waters
The aquifer partial pressures of CO2, H2S, CH4 and H2 in m a n y geothermal systems are summarized in Tables I and II and in Fig. 2. Pco~ displays a particularly good relationship with temperature, both for wetsteam and hot-water well data, to temperatures as low as 22°C. The scatter for PH~S and PH~ is much greater. Inspection o f the data points in Fig. 2 shows a tendency for PH~s and PH~ to fall on one o f two curves. One curve corresponds with all waters of ~ 200°C and waters in the range 200--300°C when containing ~ 500 ppm C1-. The other curve corresponds to all waters of > ~ 300°C and waters in the range 200--300 ° when contain-
ing > 500 ppm C1-. There is practically no overlap between data points belonging to either of the two curves so defined. According to ArnSrsson and Gunnlaugsson {1985) the two different curves for PH~S and PHi, are due to buffering of PH~S and PHi, respectively, by different mineral assemblages. Above 230°C and for "dilute" water the assemblage includes pyrite, pyrrhotite, epidote and prehnite, but pyrrhotite, epidote, prehnite and magnetite or chlorite are involved in the case o f the "saline" waters. PCH4does not show any clear dependence on aquifer temperature (Fig. 2) which would be expected at equilibrium. Examination shows that the "dilute" geothermal waters and waters of < 200°C depart considerably from equilibrium for the Fischer--Tropsch reaction: CO2 + 4H2 ~- CI-h + 2H20
(3)
Higher methane and/or lower hydrogen gas pressure would change the system towards equilibrium. By contrast, some of the "saline" waters, particularly those from New Zealand, closely approach equilibrium, a result which conforms with that o f Giggenbach (1980). Other saline waters depart considerably from equilibrium at the aquifer temperature: the Cerro Prieto waters contain less methane than would be required for equilibrium whereas the situation is the reverse for the Svartsengi and Reykjanes geothermal seawaters in Iceland. With few exceptions (some of the hottest waters) PN~ is below 2--3 bar (Fig. 2), which is somewhat higher or equal to the level anticipated if the sole source o f the nitrogen was atmospheric and dissolved in the meteoric source water (~ 0.5 umol N2 per kg H20). Above ~ 250°C Pco: is significant relative to PH~O and will thus have a significant influence at which total pressure the geothermal waters will begin to boil. The influence of Pco~ rises abruptly with temperature. At 300 C it is ~ 10 bar (abs.) and is probably as high as 30 bar (abs.) at 350°C. PH2 also significantly affects the boiling point for "dilute"
323
mo
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i o'
o~o~ e
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o
@
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o
lc;o
'
26o
'
• =- '
300
160
Aquifer temp, °C
ffi .1 ,ll
E
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oo
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o~O
o 0
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o
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~
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HOT-WATER WELLS:
•
•
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o
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e
ICELAND
i
260
3(:)0
m WAIRAKEI, KAWERAU, BROADLANDS; NEW ZEALAND
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%
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Aquifer temp. °C
WET-STEAM WELLS: ICELAND
o o
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a
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Aquifer temp. °C
08
oan~O •
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o
'
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E
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o
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~-1
•
'
Aquifer temp. °C
'
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Aquifer temp. °C JAPAN
OLKARIA; KENYA
V TONGONAN, PALINPINON; PHILIPPINES
o CERRO PRIETO; MEXICO
Fig. 2. Partial pressure for CO2, H2S, H 2, CH 4 and N2 in geothermal waters. The curve on the bottom righthand diagram representsPN~ corresponding to a concentration o f 0.5 u m o l , the quantity dissolved in meteoric water• Open symbols refer to all @aters at and above 300°C and waters in the range 200--300°C, if containing more than 500 p p m C1-. Filled symbols designate all waters below 200°C and waters in the range 200--300°C if containing less than 500 ppm C1-. The arrows in the top two diagrams o f the figure indicate the shift in the respective gas partial pressure if eq. 1, rather than eq. 2 was used to derive gas concentration in the deep reservoir water (see text)•
waters, i.e. waters containing less than 500 ppm C1- at temperatures above ~ 250°C. For "saline" waters PH2 is alSO significant above 300°C. Other gases do not exert a significant partial pressure relative to steam pressure, except N2, particularly in the case of waters from Krafla, northern Iceland. Here PN~ is as high as 40 bar (abs.). Although some o f the hottest waters contain the highest nitrogen, its activity is not predictable in terms of temperature. PH~S is below 1 bar for all waters, as well as PcH4 in the case of the "dilute" waters. For waters from Cerro Prieto (Mexico), and Kawerau and Broad-
lands (New Zealand), the methane partial pressure approaches and somewhat exceeds 1 bar (abs.). Fig. 3 shows the total pressure at which individual waters listed in Table I will begin to boil and h o w this boiling point relates to that of pure water. The following equation describes the data points in Fig. 3: l o g Ptotal = 0 . 0 0 5 4 3 T -
6 2 1 . 4 T -1
(4)
where P is in bars (abs.); and T in kelvins. In calculating the total pressure, which was done with the use of the program of ArnSrsson et al. (1982), PH~O was taken to be
324
TABLE I Gas partial pressures in geothermal reservoir water -- Data from wet-steam wells Location
1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20
Reykjanes 8, Ice. Svartsengi 4, Ice. Svartsengi 5, Ice. Svartsengi 6, Ice. Hveragerdi 2, Ice. Hveragerdi 4, Ice. Hveragerdi 6, Ice. Hveragerdi 7, Ice. Nesjavellir 5, Ice. Nesjavellir 6, Ice. ReykjabS1 1, Ice. Rey kholt 1, Ice. N~imafjall 8, Ice. N~mafjall 11, Ice. Krafla 6, Ice. Krafla 8, Ice. Krafla 9, Ice. Krafla 13, Ice. Krafla 14, Ice. Krafla 16, Ice.
21
Onuma O-6R, Japan
22 23 24
Olkaria 6, Kenya Olkaria 10, Kenya Olkaria 12, Kenya
25 26 27 28 29
Palinpinon 1, Phil. Palinpinon 7, Phil. Tongonan 202, Phil. Tongonan 404, Phil. Tongonan 407, Phil.
30 31 32 33 34 35 36 37 38
Wairakei 72, N.Z. Wairakei 81, N.Z. Kawerau 7, N.Z. Kawerau 8, N.Z. Kawerau 17, N.Z. Kawerau 19, N,Z. Broadlands 11, N.Z. Broadlands 23, N.Z. Broadlands 25, N.Z.
39 40 41 42
Cerro Cerro Cerro Cerro
79-009 79-008 79-139 79-138 79-136 79-032 81-016 81-017 72-146 83-013 79-030 79-028 79-052 81-022 81-020 81-029 79-048 81-033 81-035 81-036
CO 2
Prieto Prieto Prieto Prieto
Aquifer discharge depth .1 of:
log(partial pressure in bars abs.)
No. (for Sample Iceland No. see Fig. i)
5, Mex. 8, Mex. 35, Mex. 42, Mex.
H2S
enthalpy (J g - ' )
inflow (m)
248 *2 240 240 240 182 181 215 225 271 290 152 133 246 320 300 215 240 325 295 300
1,151 1,038 *3 1,038 *3 1,038 *3 772 *3 768 *3 921 *3 967 *3 1,190 *~ 2,062 641 *3 559 *3 1,093 2,355 1,940 921 *3 1,038 *9 1,877 2,663 1,607
1,010--1,740 1,090--1,565 835,860
CH~
0.100 - 1 . 7 5 4 0.017 - 2 . 2 4 0 -0.087 -2.319 0.079 - 1 . 9 0 7 -0.796 -1.924 -0.672 -1.955 -0.324 -1.670 -0.166 -1.504 0.841 - 0 . 5 2 9 0.610 - 0 . 0 9 5 -1.290 -3.012 - 1 . 9 4 7 -3.801 0.199 - 0 . 5 2 4 0.450 - 0 . 2 1 8 1.117 - 1 . 1 5 6 -0.111 -1.363 0.164 - 1 . 2 8 2 0.905 - 0 . 8 5 4 1.668 1.054 1.037 - 0 . 5 5 9
-1.996 -2.495 -2.237 -2.521 -1.422 1.337 -1.117 -1.169 0.502 0.867 -1 .857 -1.936 0.871 1.061 0.000 -1 .583 -0.597 0.164 0.873 0.336
-2.129 -1.446 -2.331 -1.306 -2.573 -0.695 -2.219 -0.412 -2.291 -0.815 -2.191 -0.703 -2.141 0.173 -1.943 0.107 -0.374 0.170 -1.479 -0.208 -1.900 0.491 -2.007 0.520 -0.072 0.143 -1.424 1.326 -0.762 1.104 -1.241 - 0 . 5 4 5 -1.363 -0.638 -1.382 0.464 0.590 1.576 -0.955 0.158
-0.111
-1.754
-1.996
-2.307
0.185
210
0.764 0.491 0.439
-0.362 -0.896 -0.870
0.365 0.190 0.336
-0.175 -0.764 -0.818
-0.251 -0.304 -0.418
242 261 254
2,281 2,355 2,240
0.734 1.260 0.491 1.407 1.288
-1.457 -0.845 -1.426 -0.836 -0.939
277 318 312 313 315
1,147 1,326 1,395 2,065 1,770
248 *2 244 *2 265 *2 250 *~ 261 *2 269 *2 284 *2 286 *2 298 *5
1,295 1,350 1,030 1,030 1,030 1,210 1,080 1,265 1,390
289 *2 290 *2 293 *2 281 *2
-----
0.203 - 1 . 7 1 9 -0.558 -1.828 0.699 - 1 . 2 0 6 0.878 - 1 . 1 7 4 0.650 - 1 . 2 8 4 0.947 - 0 . 9 4 8 0.961 - 1 . 4 7 7 1.404 - 1 . 0 1 4 1.353 - 1 . 1 8 7 0.694 1.071 0.941 0.771
-0.844 -0.998 -1.119 -0.867
-1 .731 -2.829 -1.759 -1.202 -1.466 -0.993 -1.454 -0.837 -0.873 -0.016 -0.352 -0 .171 -0.246
N2
temperature (°C)
H2
-0.937 -0.840 -2.838 -1.290 -0.889 -0.678 0.074 - 0 . 3 4 0 -0.059 -0.497 0.170 - 0 . 2 3 9 -0.029 -0.431 0.370 - 0 . 0 9 2 0.353 - 0 . 0 1 0 0.243 0.008 0.011 0,257
-0.759 -1.055 -0.882 -0.764
4OO 630--670 64O 405 980, 1,220 1,096 773 745 850, 1,000 1,430 1,100, 1,520 1,200 1,226 1,650 1,030, 2,030 1,050 1,078, 1,274 8O0 728, 1,025 .4 750, 855--880 2,650 2,600 1,100--1,250 1,500--1,600 1,325--1,405
• 1 More detailed data for 79 and 81 samples are given by Arn6rsson et al. (1983) and Arn6rsson and Gunnlaugsson (1985), respectively. • 2 Derived using the Na--K--Ca geothermometer. • 3 Calculated from the aquifer temperature. Boiling starts within the well. It is therefore safe to assume the discharge enthalpy to be equal to that of steam-saturated water at the respective aquifer temperature. • 4 Dominant aquifer. The data from Onuma, Japan; Olkaria, Kenya; Palinpinon and Tongonan, Philippines; the New Zealand fields; and Cerro Prieto, Mexico, were derived from the following sources: H. Sakai (pers. commun., 1982); Muna (1982); Jordan (1982); Baltasar (1980); Giggenbach (1980); a n d Nehring and d'Amore (1981), respectively.
325
TABLE II Carbon dioxide and hydrogen sulphide partial pressures in hot-water well discharges in Iceland No. (see Fig. 1 )
Sample Location No.
log PCO 2
log PH~S
Aquifer temperature
Fluid inflow (m)
(°c) 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61 62 63 64 65 66 67 68 69 70 71
79-034 79-033 79-010 79-001 79-012 79-015 81-028 79-050 79-046 81-013 81-014 79-044 81-011 81-015 79-042 79-041 79-040 81-006 79-039 81-005 79-038 81-003 79-057 79-056 79-022 79-020 79-019 79-016 79-017
Bakki 1 Oxnalaekur 1 .~rbaer 1 KlausturhSlar 1 HfisatSttir 4 Laugaland 3 Laugaland 4 Urridavatn 3 Laugar 2 Gr:~tubakki 1 Svalbardseyri 2 Laugaland 7 Gler~rgil 7 TjSrn 1 Dalv~ 10 01afsfjSrdur 3 Siglufj6rdur 7 Reykir 1 Saud~rkrSkur 1 Saud~rkrbkur 12 Varmahlfd 1 Laugarbakkar 3 ReykhSlar 2 Saelingsdalur 10 Baer 3 Seltjarnarnes 4 Reykjav~ 11 Reykir 17 Reykir 5
-2.503 -0.943 -3.397 -0.409 -3.370 -4.260 -3.754 -4.770 -5.180 -5.714 -4.893 -3.697 -4.747 -5.476 -4.804 -4.616 -4.429 -4.810 -4.427 -4.569 -3.257 -3.003 -2.860 -4.268 -2.939 -2.936 -2.921 -3.851 -3.509
-3.883 -2.947 -5.556 -2.921 -6.390 -6.857 -5.896 -7.069 -7.420 -7.914 -7.197 -6.096 -7.013 -8.021 -7.038 -7.145 -6.842 -6.951 -5.943 -6.046 -4.656 -5.037 -4.987 -5.870 -5.058 -4.503 -4.863 -6.253 -5.449
134 157 86 160 70 61 96 59 64 22 54 93 55 34 64 67 67 58 68 70 92 102 100 61 109 114 129 76 81
475--871.1 936 451 80, 120 .1 170 105 *2 750 170 250 .2, 522 *2 65 970, 1,330, 1,490 1,120, 1,440, 1,485 330, 450, 740 100 818 1,128 550 590 ? 260--471.1 83, 183 -23--285 113, 120 245, 275, 320 195--1,945 .1 927 1,750 *2 270--1,045 .1
• 1 Several aquifers were encountered in the depth range indicated. The temperature between the hottest and coldest aquifer is always less than 10°C. .2 Dominating aquifer(s). Analyses of waters collected in 1979 and 1981 are given by Arnbrsson et al. (1983) and Arnbrsson and Gunnlaugsson (1985), respectively. They also give more detailed information on physical drillhole data.
equal to that of steam-saturated water at the respective aquifer temperature. This is a satisfactory approximation when considering the scatter of data points for gas partial pressures and because PH20 is not so sensitive to total pressure for the range in question and in the presence of liquid water. Eq. 4 describes approximately the temperature--pressure relationship in boiling geothermal systems which have closely ap-
proached solute--mineral equilibrium. The equation is only approximate, since PN2 varies irregularly with temperature and because PH2S and PH2 for both "saline" and "dilute" waters are combined into one equation. Eqs. 5--9 describe the temperature dependence of Pco2, PH2Sand PH2: log Pco2
= -2.81 - 5012.7T -~ - 0.00919T + 6.464 log T
(5)
326
the range 200--300°C if they contain more than 500 p p m C1-. Eqs. 7 and 9 (with subscripts d) are valid for all waters below 200°C and for waters in the range 200--300°C if they contain ~ 500 ppm C1-. Eq. 5 describing the temperature dependence of Pco~ differs somewhat from that given b y Giggenbach (1981), b u t he used data largely in the range 200--300°C. Eq. and his equation yield 0.35 and 0.05 bar (abs.) at 200°C, practically the same at 250°C, but the difference at 300°C is 27 bar (abs.). Visual inspection of the data points for Pco~ in Fig. 2 above 200°C does not reflect an unambiguous relation with aquifer temperature, which has recently been a source of disagreement (Giggenbach, 1982; Grant, 1982). However, b y considering waters over a larger range of temperatures, as is the case for the present study (see Fig. 2), such dependence is evident.
50 .Q
XI
u~ u) L. Q.
~o
!
150
KRAFLA WELL 14
4. Boiling
200 300 Temperature °C
Fig. 3. Relation between total pressure (steam and gases) and aquifer temperature in geothermal waters. The broken line inducates temperature---pressure relationship for water from well 14 in Krafla during adiabatic boiling and equilibrium degassing.
log Pg~s(s) = - 1 8 . 7 5 - 24738.9T -1 - 0 . 1 0 1 3 3 T + 43.170 log T
(6)
log PH~S¢d) = - - 2 . 7 6 - 5 7 5 8 . 2 T -1 - 0.00850T + 6.359 log T logPH~(s )
=-10.14-
1ogPH~(d)
(7)
1 7 7 6 3 . 1 T -~
0 . 0 4 0 8 4 T + 23.368 log T
(8)
= 9.85 + 7290.3T -1 + 0 . 0 7 2 0 2 T -
22.651 log T
(9)
where P is again in bars (abs.); and T in kelvins. Eqs. 6 and 8 {with subscripts s) are valid for all waters of > 300°C and for waters in
point
curves
Boiling of initially equilibrated geothermal water in a closed system will eventually lead to almost quantitative transfer of the dissolved gases into the steam and, as a result, the gas partial pressures will be drastically reduced. This implies that a large pressure drop will accompany a small temperature drop during the early stages of boiling, if gas partial pressures are significant relative to the vapour pressure. Under closed-system conditions each geothermal water will possess a unique boiling point curve dictated b y initial temperature and its gas composition. One such curve is shown in Fig. 3 for well 14 in the Krafla geothermal field and in Fig. 4 for other selected wells in terms of drop in total gas pressure for adiabatic boiling. The program of Arn6rsson et al. (1982) was used to calculate these curves, simultaneously taking into account the distribution of eight gases (CO2, H2S, H2, CH4, N2, 02, Ar and NH3) between the water and steam phases. The initial slope of the boiling point curves will depend on which gas dominates
327
the pressure, being steepest when the least soluble gases, H2 and N2, contribute most to the total gas pressure. As the boiled water becomes more and more depleted in total gas with continued steam formation all boiling point curves approach more and more closely the boiling point curve for pure water. Degassing accompanying boiling in geothermal systems m a y often be incomplete. However, this has little effect on the shape of the boiling point curves, even if the degas1 sing is as little as ~ of m a x i m u m (that is equilibrium degassing). The reason is that all the gases are strongly partitioned into the steam phase. Salinity, which will reduce gas solubility, also has little effect on the shape of the boiling point curve. i
Krafla well 16
101 .,0
Hveragerdi well 7
,//
/
..D
ffl In
Je It:e""
~. 10-1 O)
10-3
;/ I
I
200
I
I
300 Temperature °C
Fig. 4. Reduction in gas pressure for selected geothermal waters during adiabatic boiling and equilibrium degassing.
5. Discussion
Most of the data considered in the present study provide reliable data on gas partial pressures, namely data from hot-water and
wet-steam wells where the boiling starts within the well. The data from "high-" enthalpy wells might be regarded as suspect and the gas content of the well discharge may not correspond with that dissolved in the water at depth, as indicated b y eq. 2. When comparing these two types of data it is seen that they give a coherent picture of the relation between gas partial pressure and aquifer temperatures (Fig. 2). From this it is concluded that the model adopted and described in a previous section is satisfactory, as well as the temperature dependence of gas partial pressures expressed by eqs. 5--9. It is evident that the same mineral assemblages will not buffer the gas partial pressures over the temperature range of 22--325°C considered in this study. Yet a smooth temperature dependence is observed for Pco~, the reason for this most likely being that the free energy difference between the different assemblages is small. Such is often known to be the case for metamorphic mineral assemblages. The same does not, however, hold for PH2S and PHi, as the different mineral buffers fix the respective gas pressures at significantly different levels at any particular temperature. It has not been established which minerals constitute these buffers at low temperatures, but at elevated temperatures (above ~ 200--250°C) Giggenbach {1981) and Arn6rsson and Gunnlaugsson (1985) have evaluated which assemblages are most likely to be involved. It is often observed that temperature profiles in geothermal wells deviate from the boiling point curve of pure water to lower temperatures for the measured total pressure. This deviation may in some instances, particularly at temperatures above 300°C, be due to the effect of gas pressure on the boiling point. In other wells slow heating-up may eventually bring temperatures to the boiling point for pure water. Here the reason m a y be that boiling and degassing occurs in the closed well, as evidenced b y accumulation of gas under the wellhead. The flow m a y either be provoked by condensation o f steam at
328 high levels in t h e well or b y f l o w o f fluid into a s h a l l o w e r aquifer. With t h e aid o f eqs. 5 - - 9 , analyses o f t h e c o m p o s i t i o n o f t h e gas u n d e r the wellhead and direct measurement of total gas p r e s s u r e ( t o t a l p r e s s u r e m i n u s s t e a m pressure as i n d i c a t e d b y t e m p e r a t u r e ) it m a y be possible t o e v a l u a t e t h e t e m p e r a t u r e o f t h e d o m i n a t i n g a q u i f e r in t h e well.
0° to 370°C. Geochim. Cosmochim. Acta, 46: 1513--1532. Arn6rsson, S., Gunnlaugsson, E. and Svavarsson, H., 1983. The chemistry of geothermal waters in Iceland, II. Mineral solubilities a n d independent variables controlling water composition. Geochim. Cosmochim. Acta, 47: 1513--1532. Baltasar, A.J., 1980. Interpretation of the water and gas chemistry from three geothermal areas in the Philippines Manito in Albany, Biliran Island and Tongonan in Leyte. U.N. Univ. Geotherm. Train. Prog., Reykjav~, Rep. 1980-3, 55 pp. Giggenbach, W.F., 1980. Geothermal gas equilibria. Geochim. Cosmochim. Acta, 44: 2021--2032. Giggenbach, W.F., 1981. Geothermal mineral equilibria. Geochim. Cosmochim. Acta, 45: 393--410. Giggenbach, W.F., 1982. Reply to a Comment by M.A. Grant. Geothermal mineral equilibria. Geochim. Cosmochim. Acta, 46: 2681--2684. Grant, M.A., 1982. On the lack of a unique relation between CO s partial pressure and temperature in geothermal systems. Comment on Geothermal mineral equilibria, by W.F. Giggenbach. Geochim. Cosmochim. Acta, 46: 2677--2680. Jordan, O.T., 1982. Implication of solution--mineral equilibria on the exploitation of the S-Negros geothermal field, Philippines. U.N. Univ. Geotherm. Train. Prog., Reykjav~, Rep. 1982-7, 67 pp. Muna, Z.W., 1982. Chemistry of well discharges in the Olkaria geothermal field, Kenya. U.N. Univ. Geotherm. Train. Prog., Reykjav~, Rep. 1982-8, 38 pp. Naumov, G.B., Ryzhenko, B.N. and Khodakovsky, I.L., 1971. Handbook of thermodynamic data. U.S. Geol. Surv., Rep. WRD-74-001 (translated of a Russian report). Nehring, N.L. and d'Amore, F., 1981. Gas chemistry and thermometry of the Cerro Prieto geothermal field. Proc. 3rd Syrup. on the Cerro Prieto Geothermal Field, Baja California, March 1981. Steffinsson, V. and Bj6rnsson, S., 1982. Physical aspects of hydrothermal systems. In: Continental and Oceanic Rifts. Geodyn. Ser., 8: 123--145. -
Acknowledgements T h i s article was originally p r e s e n t e d at the Fourth International Symposium on W a t e r - - R o c k I n t e r a c t i o n . I t is n o w p r e s e n t e d in an e x t e n d e d f o r m b y i n c o r p o r a t i n g d a t a f r o m g e o t h e r m a l fields, n o t o n l y f r o m Iceland, b u t f r o m o t h e r c o u n t r i e s as well. Dr. Einar Gunnlaugsson from the Reykjav~ D i s t r i c t H e a t i n g Service is specially t h a n k e d f o r his assistance w i t h c o m p u t e r w o r k and Dr. V a l g a r d u r Steffinsson o f t h e N a t i o n a l E n e r g y A u t h o r i t y , Reykjaw~k, f o r p r o v i d i n g s o m e o f t h e p h y s i c a l drillhole data. P r o f e s s o r H i t o s h i Sakai is also t h a n k e d f o r p r o v i d i n g c h e m i c a l a n d p h y s i c a l drillhole d a t a f o r t h e O n u m a field, J a p a n .
References Arn6rsson, S. and Gunnlaugsson, E., 1985. New gas geothermometers for geothermal exploration -calibration and application. Geochim. Cosmochim. Acta, 49 (in press). Arn6rsson, S., Sigurdsson, S. and Svavarsson, H., 1982. The chemistry of geothermal waters in Iceland, I. Calculation of aqueous speciation from
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