Fluid-Fluid Interactions in Geothermal Systems

7/22/2019 Fluid-Fluid Interactions in Geothermal Systems

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9Reviews in Mineralogy & Geochemistry

Vol. 65, pp. 259-312, 2007

Copyright Mineralogical Society of America

1529-6466/07/0065-0009$10.00 DOI: 10.2138/rmg.2007.65.9

Fluid-Fluid Interactions in Geothermal Systems

Sten Arnrsson, Andri StenssonInstitute of Earth Sciences, University of Iceland

Sturlugata 7, IS-101 Reykjavk, [email protected], [email protected]

Jn rn BjarnasonIceland GeoSurvey

Grensasvegur 9, IS-108 Reykjavik, [email protected]

INTRODUCTION

The main goal o geothermal geochemistry research is to identiy the origin o geothermalfuids and to quantiy the processes that govern their compositions and the associated chemicaland mineralogical transormations o the rocks with which the fuids interact. The subjecthas a strong applied component: Geothermal chemistry constitutes an important tool or theexploration o geothermal resources and in assessing the production characteristics o drilledgeothermal reservoirs and their response to production. Geothermal fuids are also o interest asanalogues to ore-orming fuids. Understanding o chemical processes within active geothermalsystems has been advanced by thermodynamic and kinetic experiments and numerical modelingo fuid fow. Deep drillings or geothermal energy have provided important inormation onthe sources and composition o geothermal fuids, their reaction with rock-orming minerals,migration o the fuids, and fuid phase separation and fuid mixing processes.

Based on ndings to date, geothermal fuids may be classied as primary or secondary.Primary fuids are those ound in the roots o geothermal systems. They may constitute a mix-ture o two or more fuids, such as water o meteoric origin, seawater and magmatic volatiles.Several processes can lead to the ormation o secondary fuids, such as the boiling o aprimary fuid that separates it into liquid and vapor and the un-mixing o a very hot brine byits depressurization and cooling. Further, secondary geothermal fuids orm by the mixing odeep fuids with shallow ground water or surace water. In this chapter we summarize the geo-hydrological and geochemical eatures o geothermal systems and delineate the processes thatproduce the observed chemical composition o the various types o geothermal fuids ound inthese systems. The main emphasis is, however, on gas chemistry and the assessment o fuidphase separation below hot springs and around discharging wells drilled into liquid-dominatedvolcanic geothermal systems.

BASIC FEATURES OF GEOTHERMAL SYSTEMS

Geothermal systems consist o a body o hot rock and hot fuid, or hot rock alone, in aparticular rock-hydrological situation. The lietime o individual systems is generally poorlyknown, as are fuid circulation times. Estimated ages o geothermal systems lie in the range osome 0.1 to 1 million years (Grindley 1965; Browne 1979; Arehart et al. 2002). Fluid circulationtimes may range rom ew hundreds o years, or even less, to more than 10,000 years (Arnrsson1995a; Sveinbjrnsdttir et al. 2001; Mariner et al. 2006).


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260 Arnrsson, Stefnsson, Bjarnason

The fuids o geothermal systems are essentially meteoric water or seawater by origin ormixtures thereo. Some geothermal fuids contain a signicant component o magmatic volatiles.

Mixing o these source fuids and their interaction with the rock produces several types o geo-thermal fuids, as exemplied in Figure 1 or a system associated with an andesitic volcano.

The maximum depth o meteoric or seawater circulation is the brittle/ductile transitiondepth o rocks (Fournier 1991). This transition has been inerred rom the depth o earthquakeoci, requently reaching depths o 5 to 8 km (Klein et al. 1977; Bibby et al. 1995) wheretemperatures may be 400 C or higher (e.g., Kissling and Weir 2005). Fluid circulation ingeothermal systems is essentially density driven when temperatures at the base o thecirculation are above ~150 C (see also Driesner and Geiger 2007, this volume, or parameterscontrolling hydrothermal convection).

The ultimate source o heat to geothermal systems is decaying radioactive elements, par-ticularly U, Th and K, and the Earths gravity eld. Rising magma transports heat to the uppercrust, and convecting fuids in permeable rocks transport the heat to even shallower depths,creating a geothermal system.

Types o geothermal systems

Several classication schemes have been proposed or geothermal systems. The mostimportant ones are high-temperature and low-temperature systems (Bdvarsson 1961), hot-water (liquid-dominated) and vapor-dominated systems (White et al. 1971) and volcanic andnon-volcanic systems (Go and Janik 2000). High-temperature systems are generally volcanicand low-temperature systems non-volcanic.

Volcanic geothermal systems typically occur in areas o active volcanism where permeabilityand geothermal gradient are high. The heat source may be a major magma intrusion, a sheeted

Degassing

Coldwater

Two-phase zonewater+steam

Magma

Conductive boundar

y

Neutral pHNa-Cl water

Acid SO -Cl water

Brine4

Figure 1. Schematic section o a volcanic geothermal system depicting theorigin o dierent primary fuid types.


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Fluid-Fluid Interaction in Geothermal Systems 261

dyke complex, or a complex o minor intrusions. In liquid-dominated systems, the fuid isdominantly liquid water by volume (White et al. 1971). In vapor-dominated systems, vapor at

~240 C lls permeable channels, although liquid water occupies inter-granular pore spaces(Fig. 2). Very hot saline liquid (>400 C) is known to exist below the vapor-dominated zone(Fournier 1991; Barelli et al. 1995; Gianelli and Ruggieri 2002). Vapor-dominated systems areconsidered to evolve by the boiling down o earlier liquid-dominated systems, i.e., the supplyo heat to the system in relation to fuid circulation rates is sucient to vaporize a substantialraction o the circulating fuid.

Geological structure o volcanic geothermal systems

Structurally, no two volcanic geothermal systems are identical. To describe the resultso extensive studies o a particular system requires a model, which is unique to that system(e.g., Hochstein and Browne 2000). However, many geothermal systems have some commoneatures, which are the geological structure and the nature o the heat source.

The geological structures o volcanic geothermal systems on diverging and convergingplate boundaries dier in some respects. On converging plate boundaries, they are requentlyassociated with andesitic volcanoes, and the heat source may be magma oshoots rom themain magma source or a deeper magma body in the roots o the volcano. On diverging plateboundaries, the heat source is usually a sheeted dyke complex. Volcanic geothermal systemsare commonly associated with ring structures or calderas. This structure is the consequenceo rapid emptying o a shallow magma body. The magma at converging plate boundaries ischaracteristically high in volatile components relative to the basaltic magma generated be-low diverging plate boundaries. This has signicant consequences or the composition o thegeothermal fuids in these two types o geological settings. The chemical composition andmineralogy o the host rocks also have major infuence on the chemical composition o thegeothermal fuid.

Temperature and pressure

Measured downhole temperatures in drilled geothermal systems range rom ambient toover 400 C. Depending on P and Tthe fuids o geothermal systems may be single phase or

two-phase. Vapor-dominated systems represent a special type o the latter. Their mobile phase

TemperaturePressure

Dep

th

TP

Cap rock

Boiling water

Steam

Water

Figure 2. Schematic model o hydrological, fuid, temperature and pressureconditions in a vapor-dominated geothermal system. Based on Fournier (1981).


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deposits (Buat-Menard and Arnold 1978; Lepel et al. 1978; Le Cloarec et al. 1992; Symondset al. 1992; Hedenquist and Lowenstern 1994; Hedenquist et al. 1994).

GEOTHERMAL FLUIDS

Primary and secondary geothermal fuids

In this contribution, geothermal fuids at the bottom o the convection cell (base-depth)are termed primary geothermal fuids. They may be a mixture o two or more fuid componentssuch as meteoric and seawater and magmatic volatiles. The main types o primary fuids areNa-Cl waters, acid-sulate waters and high salinity brines. When primary fuids rise towards thesurace, they can undergo fuid phase separation and fuid mixing to orm secondary geothermalfuids. The most important processes that lead to the ormation o secondary geothermal fuidsare:

1) Depressurization boiling to yield boiled water and a steam phase with gas.

2) Phase separation o saline fuids into a hypersaline brine and a more dilute vapor.

3) Vapor condensation in shallow ground water or surace water to produce acid-sulate,carbon-dioxide or sodium bicarbonate waters.

4) Mixing o CO2 gas rom a deep source with thermal ground water.

5) Mixing o geothermal fuids with shallower and cooler ground water.

Chemical composition o primary fuids

The chemical composition o primary geothermal fuids is determined by the compositiono the source fuids and reactions involving both dissolution o primary rock-orming mineralsand deposition o secondary minerals, as well as by adsorption and desorption processes. The

18

16

14

12

10

8

6

4

2

00 200 400 600 800 1000 1200

Temperature (C)

Dep

th(m)

Molten lava

Conductive boundary

Boiling

Figure 3. Temperature prole in a hole drilled into molten lava fow on the island o Heimaey o the southcoast o Iceland and a heat-transer model or volcanic geothermal systems. Water was pumped onto thelava. The water percolated to a depth o 4 m where it was all converted into steam. The rising steam keptthe lava above 4 m depth at 100 C. The temperature o the molten lava was ~1000 C. Heat was transerredconductively through the layer at 4 to 6 m depth to the circulating water. Continued cooling involvesdownward migration o the conductive layer.


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source fuids are usually meteoric water or seawater or a mixture thereo. Components oconnate, magmatic and metamorphic fuids may also be present in geothermal fuids.

Na-Cl waters. The dissolved salt in Na-Cl waters is mainly NaCl. This type o water isthe most common in geothermal systems. Chloride concentrations typically range rom onlya ew hundred to a ew thousand ppm. They are lowest in waters hosted in basaltic rocks(Sigvaldason and skarsson 1976; Arnrsson and Andrsdttir 1995) but highest in fuidswhich have interacted with sedimentary rocks containing evaporites (Helgeson 1968; White1968; Table 1). The salinity o geothermal fuids is determined by the availability o solublesalts. These salts may be leached rom the aquier rock or added to the geothermal fuid by deepmagmatic fuids. Alternatively, saline fuids may orm through reactions between magmaticHCl and rock-orming minerals.

The concentrations o most major elements in Na-Cl waters are xed by close approachto local equilibrium with secondary minerals i temperatures are above ~100 to 150C (e.g.,Browne and Ellis 1970; Giggenbach 1980, 1981; Arnrsson et al. 1983; Hedenquist 1990;Gudmundsson and Arnrsson 2005; Karingithi pers. comm.). The only conservative major

component in these waters is Cl. The mineral-solution equilibria constrain ion activity ratiosand the activities o neutral aqueous species other than Cl-bearing species, including reactivegases like CO2, H2S and H2,which may be largely o magmatic origin. Some systems closelyapproach redox equilibrium (Seward 1974) while others signicantly depart rom it (Stenssonand Arnrsson 2002; Stensson et al. 2005).

The concentrations o many trace elements (e.g., Ag, Fe, Cu, Pb, Zn) in Na-Cl geothermalwaters are clearly controlled by sulde mineral deposition (e.g., Simmons and Browne 2000;Reyes et al. 2002). These elements typically orm cations in solution. Trace elements that ormlarge simple anions or oxy-anions in solution may have high mobility and even show incompat-ible behavior (e.g., Br, I, As, Mo, W) (Arnrsson 2003; Arnrsson and skarsson 2007).

Acid-sulfate waters. Deep acid-sulate fuids have been encountered in many volcanicgeothermal systems, particularly in association with andesitic volcanoes (Truesdell 1991;Giggenbach and Corrales 1992; Sanchez and Arnrsson 1995; Kiyota et al. 1996; Salonga1996; Akaku et al. 1997; Hermoso et al. 1998; Gherardi et al. 2002; Matsuda et al. 2005; Tello etal. 2005). The acidity is caused by HCl or HSO4 or both, and evidence indicates that it mostlyorms by transer o HCl and SO2 rom the magmatic heat source to the circulating fuid.

When measured at 25 C, the pH o fashed acid-sulate water collected at the wellheadmay be as low as 2. The pH o the water is near neutral at the high temperature in the aquier,however. Production o acidity upon cooling is related to the increased acid strength o HSO4with decreasing temperature (Fig. 4).

The most important dierence between the Na-Cl and acid-sulate waters is that the mainpH-buer o the ormer is CO2/HCO3, but HSO4/SO42 in the latter. Compared to Na-Cl waters,acid SO4-Cl waters contain higher concentrations o SO4 and some minor elements, such as Feand Mg, which are contained in minerals with pH-dependent solubility.

Elevated Cl concentrations (up to 120 ppm by weight) have been measured in superheatedvapor at The Geysers and Larderello vapor-dominated elds (DAmore et al. 1977; Haizlipand Truesdell 1988) and the Krafa liquid-dominated eld (Truesdell et al. 1989). The Cl inthe vapor is transported as HCl. A high Cl concentration in the vapor is due to evaporation obrine. The Cl concentration o the vapor is, however, also aected by the pH o the brine andthe temperature o separation o vapor and brine (Fig. 5).

High salinity waters.Geothermal brines can orm in several ways. Brine-orming processesinclude dissolution o evaporites by water o meteoric origin (Helgeson 1968; White 1968) andreaction between some primary minerals o volcanic rocks and magmatic HCl. Connate hot-


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Table1.Chemicalanalysesothemajortypesogeothermalwaters.Concentrationsareinmg/L.

1

2

3

4

5

6

7

8

9

10

11

12

h(kJ/kg)a

1256

1376

1257

1423

1287

995

1315

1839

1653

2306

1850

TCb

237

260

251

248

292

278

248

284

251

262

280

260

Sampl.pressurec

0

25.0

0

0

0

0

0

0

0

0

9.8

25.3

pH/Cd

4.8/25

5.16/22

6.34/25

8.14/2

5

8.13/25

3.11/25

7.70/15

8.10/20

9.07/25

8.67/25

7.46/20

8.35/21

SiO2

550.0

668.8

660.0

361.0

741.0

910.0

557.0

784.0

643.0

855.0

815.0

547.0

B

6.72

159.0

30.0

39.9

21.0

26.2

52.0

6.48

6.80

1.73

2.32

Na

38839

9903.0

5105.0

1764.0

860.0

3117.0

1256.0

910.0

557.0

1283.0

214.4

100.5

K

6250

1314.0

994.0

277.0

175.4

950.0

200.0

155.0

92.0

208.0

43.8

15.6

Mg

34.0

1.05

0.140

0.20

0

0.190

25.00

0.020

0.010

0.010

0.070

0.049

0.002

Ca

20630

1548.0

285.0

68.4

10.80

82.0

26.7

1.3

0.73

0.66

0.80

0.40

Fe

4.4

0.325

0.180

0.060

282.0

0.005

0.010

0.020

0.020

0.027

0.032

Al

0.092

1.090

0.310

0.660

0.670

1.080

1.646

CO2e

54.8

47.4

14.6

27.7

51.2

0.0

64.1

191.0

74.0

2465.0

200.6

3.5

SO4

12.0

24.3

20.5

137.6

27.0

508.0

34.2

30.0

28.0

112.0

8.6

20.2

H2S

3.4

3.6

0.0

1.6

0.0

1.02

3.96

53.5

97.8

F

4.0

0.14

3.6

6.9

6.1

69.0

105.0

1.36

0.80

Cl

103000

20534

9074.0

3026.0

1377.0

6175.0

2183.0

1414.0

764.0

240.0

113.1

28.7

aDischargeentha

lpy.baquiertemperature.cinbar-g.dpH/temperature

omeasurement.eTotalcarbonatecarbonasCO2.

1:Na-Ca-Clbrine,AsalDijbuti(DAmoreetal.1998)

2:Seawatergeo

thermalsystem,Reykjanes,well15,Iceland(Giroud,pers.comm.)

3:Andesiticvol

canicgeothermalsystem,Tongonan,well510,Philippines(Angcoy,pers.comm.)

4:Andesiticvol

canicgeothermalsystem,Momotombo,well2,Nicaragua(Arnrsson1997)

5:Volcanicsysteminsilicic-andesiticrocks,Zunil,well3,Guatemala

(Arnrsson1995b)

6:Acidicsulphate-chloridewater,Mahanagdong,well9,Leyte,Philippines(Angcoy,pers.comm.)

7:Volcanicgeothermalsysteminsilicicvolcanics,Wairakei,well24,NewZealand(Mahon,pers.comm.)

8:Volcanicsysteminsilicic-andesiticrocks,Broadlands,well28,New

Zealand(Mahon,pers.comm.)

9:Volcanicgeothermalsystemsinbasalttotrachytevolcanics,Olkaria

,well2,Kenya(Karingithipers.comm)

10:Volcanicgeothermalsystemsinbasalttotrachytevolcanics,Olkaria

,well301,Kenya(Karingithipers.comm)

11:Volcanicgeothermalsysteminbasalt,Krafa,well20,Iceland(GudmundssonandArnrsson2002)

12:Volcanicgeothermalsysteminbasalt,Nmajall,well11,Iceland(G

udmundssonandArnrsson2002)


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water brines have been encountered in sedimentary basins (White 1965). Brines may orm byfuid phase separation through cooling and depressurization o moderately saline geothermalfuids in which case they are secondary.

A large region o fuid immiscibility exists in P-T-Xspace or saline geothermal fuids ohigh temperature (Heinrich et al. 2004; see also Liebscher 2007). The crest o the immiscibilitysurace (the critical curve) extends rom ~25 wt% NaCl at ~700 C and 1250 bar (~5 km depthor lithostatic pressure) to the critical point o pure water at 374 C and 221 bar. In volcanicgeothermal systems containing saline fuid, the temperature may be too high and the pressuretoo low to reach the immiscibility surace, in which case a two-phase geothermal fuid existsto the base o the system.

-7

-6

-5

-4

-3

-2

-1

0

0 50 100 150 200 250 300 350

logK

Temperature (C)

HSO4-

= H+ + SO4-2

Figure 4. The temperature depen-dence o the dissociation constant othe bisulate ion. The acid strengtho HSO4 decreases rapidly with ris-ing temperature, or by 4 orders omagnitude when going rom 25 to300 C. (Based on thermodynamicdata in Naumov et al. 1971).

-3

-2

-1

0

1

2

3

4

0 10000 20000 30000 40000

logHClinvapor(ppm)

Cl in fluid (ppm)

200

220

240

260280

300

320

340

Figure 5. The concentrations o HClin steam at selected temperatures(C), as indicated, as a unction o theCl concentration in the water withwhich the steam has equilibrated.The pH o the water was selected as5. The results are based on values orthe association constant or HCl asgiven by Ruaya and Seward (1987)and the distribution coecient orHCl between liquid water and vaporaccording to Simonson and Palmer(1993).


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Many metals (Ag, Au, Cu, Mo, Pb, Sn, W, Zn) orm complexes with Cl, HS and OH atmagmatic temperatures that partition strongly into the magmatic fuid during crystallization

(Hedenquist and Lowenstern 1994). As this fuid escapes rom the melt into the country rock,these metals together with magmatic gases are transported into the geothermal fuid. Mixingo the magmatic and geothermal fuids and their subsequent interaction with rock-ormingminerals leads to brine ormation, i the magma is rich in HCl. Cooling and transormation omagmatic SO2 into H2S leads to precipitation o metallic suldes. Porphyry ore-deposits areconsidered to orm in this way (Hedenquist and Lowenstern 1994; Fournier 1999; see alsoHeinrich 2007a, this volume).

Secondary fuids

Steam-heated acid sulfate waters. In many high-temperature geothermal elds, suracemaniestations consist mostly o steam vents (umaroles), steam-heated surace water andhot intensely altered ground (Fig. 6). Condensation o H2S-bearing steam by heat loss ormixing with surace water and oxidation o the H2S leads to the ormation o native sulur,thiosulate, various polysuldes and ultimately sulate (Xu et al. 1998, 2000; Druschel et al.

2003). Steam-heated acid-sulate waters are characterized by low Cl and relatively high sulateconcentrations (Table 2). It is not uncommon that the pH is


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Table 2. Major and trace element composition o the main types o secondary geothermal waters.Also shown (6-8) is the chemical composition o primary geothermal waters in silicic and basaltic

rocks, that are meteoric and seawater by origin, respectively. Concentrations are in mg/L ormajor components but in g/L or trace components.

1 2 3 4 5 6 7 8

Discharge enthalpya 1167 1256TC 100 94 60 29.4 65 93.0 238 260Sampling pressureb 18.5 25.0pH/Cc 2.55/22 1.17 6.21/25 6.32 7.98/26 8.61/20 8.80/18 5.16/22

Major components

SiO2 226.0 182.0 212.0 67.0 176.2 243.0 500.0 668.8Na 27.0 535.0 417.7 184.0 209.7 331.0 147.0 9903.0K 2.44 78.3 30.6 70.0 35.9 9.45 18.0 1314.0Mg 68.2 15.6 27.2 274.0 1.0 0.001 0.0016 1.05Ca 94.9 56.1 99.7 727.0 10.0 1.0 2.44 1548.0

CO2d

0.0 1145.0 2575.0 300.6 SO4 1363.0 1598.0 41.3 1078.0 19.9 20.7 290.8 24.3H2S


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by its salinity and gas content. Increasing salinity raises the boiling point (Bischo and Pitzer1989), whereas dissolved gases lower it. The critical point o pure water is at 374 C and 221

bars. The corresponding numbers or seawater are 405 C and 302 bars. Boiling o a risinggeothermal liquid starts at a depth where the sum o the water vapor pressure and all dissolvedgas partial pressures become equal to the hydrostatic pressure.

The boiling point curve

The solid line in Figure 8 shows how the boiling point o pure water changes with depth.In this gure, pressure has been converted into depth (length o a water column) assumingthat water at the surace is at 100 C. The increase in temperature and pressure with depthcorresponds to a column o liquid water that is at the boiling point at all depths. The curveo Bjrnsson (1966) has been updated here using the IAPWS Industrial Formulation 1997or the Thermodynamic Properties o Water and Steam. The slim curve in Figure 8 showshow the boiling point or a 3.2 wt% NaCl solution varies with depth. The salt-water curveis based on the experimental data o Bischo (1984) and Bischo and Rosenbauer (1984,1985) on temperature, pressure and specic volume o a 3.2 wt% NaCl solution along the

liquid-vapor phase boundary. According to Bischo and Rosenbauer (1984), a 3.2 wt% NaClsolution is a good proxy or seawater. The boiling point curves or pure water and seawaterare quite similar at temperatures below about 300 C but they diverge considerably at highertemperatures. The critical point o seawater is at about 405 C and 302 bar. This point occursat about 5200 m depth in a seawater column that reaches the surace and is at the boilingpoint at all depths. By comparison, the critical point o pure water would be ound at a deptho only 3500 m. The results shown in Figure 8 and those presented by Heinrich et al. (2004)show that the critical temperature and critical pressure (depth) o geothermal fuids varyenormously with salinity.

The ollowing equations describe satisactorily the boiling point with depth or pure waterand a 3.2 wt% NaCl solution (seawater), respectively:

D T T Tw = + + + 0 01120276 7 016900 10 2 395989 10 21 2092 5 3 23 10. . . . 111

0 01091091 6 614698 10 2 165155 102 5 3

= + + log( )

. . .

T

D T Ts

(1) + 23 10 42 04005 2T T. log( ) ( )

6000

5000

4000

3000

2000

1000

0

100 200 300 400

Depth(meters)

Temperature (C)

405C302 bars

374C221 bars

Figure 8. Boiling point curves withdepth or pure water (solid line) and3.2 wt% NaCl solution (thin line).A 3.2 wt% NaCl solution is a goodproxy or seawater. The curve orpure water is based on the IAPWSIndustrial Formulation 1997 or thethermodynamic properties o waterand steam, but that or seawater wasretrieved rom the experimental datao Bischo and Rosenbauer 1985)on temperature, pressure and specicvolume or a 3.2% NaCl solution alongthe liquid-vapor boundary.


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Dw andDs denote the depth in m or pure water and a 3.2 wt% NaCl solution, respectively, andTthe temperature in C.

Eect o dissolved gases

The gas content o an aqueous fuid aects its boiling point. A vapor phase begins toorm when the sum o the water vapor pressure and the partial pressures o all dissolved gasesbecomes equal to the hydrostatic pressure (Ph)

P P Ph ss

= + (3)H O2

where PH2O represents the water vapor pressure and Ps the partial pressure o gas s. Thestandard molal solubility constant (Henrys Law Constant) o gas s (Ks) in aqueous solutionis given by

a K fsl

s= (4)s

where Ks is in moles kg1 bar1 and asl designates the activity o the gas in liquid water andfs its

ugacity above the solution. The relationships between the activity (asl

) and concentration (msl

)o gas s on one hand and the ugacity and partial pressure on the other are given by

a msl

s s

l= (5)

and

f Ps s s= (6)

Here, s is the activity coecient o dissolved gas s and s the ugacity coecient. The value os can be obtained rom the dimensionless compressibility actorZs and the modied Redlich-Kwong equation (see e.g., Nordstrom and Munoz 1994):

ln ( ) ln ln.

s ss R K R K

R K

R K

s

ZZ b P

T

a

T b

b P

Z T=

+

R R

R- -

-

-1 11 5

(7)

Here,Zs is equal to PV/RT, where Vis the molar volume o a real gas, P is pressure, R the gasconstant and Tabsolute temperature. The constants aR-Kand bR-Kare equal to 0.4275R2Tc2.5/Pcand 0.0866RTc/Pc, respectively (Nordstrom and Munoz 1994), where Tc and Pc are the criticaltemperature and pressure, respectively. Nordstrom and Munoz (1994) list values o Tc and Pcor some common gases (Table 3). They also provide data or Zc, the critical compressibilityactor. WithTc, Pc and Zc, known,Zs can be obtained rom the ollowing equation:

Z P V

T

s

Z

P V

T(8)

c

c c

c

=( / )( / )

/

The activity coecient, s in Equation (5), can be obtained experimentally as describedby Nordstrom and Munoz (1994) and Anderson (2005). It can also be obtained rom theSetchnow equation

logs s Bk m= + (9)

where ks is the Setchnow coecient and mB the molar concentration o neutral salt B insolution. As an approximation, mBin Equation (9) may be taken to be 0.1 (Anderson 2005).The Pitzer equations (see Anderson 2005) provide a more detailed and accurate treatment ouncharged species than Equation (9). For neutral species, the extended Debye-Hckel equationo Helgeson (1969) reduces to

log = (10)s B I


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whereIrepresents the true ionic strength o the solution, i.e., the ionic strength ater correctionhas been made or ion pairing and complexing. Helgeson et al. (1981) updated Equation (10)or neutral species by replacingB with revised values.

For dilute solutions, it is a reasonable approximation to take s =1, i.e., asl = ms

l. This is,however, not the case or solutions o high ionic strength. For such solutions, Equation (5)needs to be solved using either (9) or (10). At pressures existing in most drilled geothermalsystems,Zs can be taken to be equal to unity or common gases, i.e., the gases can be taken tobehave ideally.

Gas pressures in geothermal systems are highly variable, ranging rom less than 1 bar,such as at Wairakei in New Zealand and Svartsengi in Iceland to as much as ~100 bar atBroadlands in New Zealand and Olkaria in Kenya. The partial pressures o individual gases

in the geothermal systems may be determined by their supply to the geothermal fuid or xedby specic temperature-dependent mineral-gas or gas-gas equilibria (e.g., Gudmundsson andArnrsson 2002; Karingithi pers. comm.).

Liquid-vapor separation under natural conditions

Boiling under natural conditions in up-fow zones o geothermal systems is maintained bydepressurization o the rising hot fuid. The boiled fuid may emerge at the surace in springsor fow laterally rom the up-fow zone underground, depending on the depth o the groundwater table o the boiled fuid.

Vapor rising above the ground water table may condense partly or totally by conductiveheat loss or by mixing with water in perched aquiers or surace water. In this way, vapor-heated secondary geothermal fuids orm. Most o the gases initially present in the deep aquierfuid will be transerred into the vapor. They may dissolve partly in water o perched aquiersand partly rise to the surace. Dissolution o CO2 and H2S in such water may turn it acid and

reactive. The H2S may be oxidized or precipitated as suldes. Carbon dioxide is transerredinto bicarbonate through reaction with the rock and may precipitate to some extent as calciumcarbonate.

Vapor-dominated systems

In vapor-dominated geothermal systems, vapor is the continuous phase in ractures,whereas liquid water lls partially or totally intergranular pore spaces (see Fig. 2). Temperature

Table 3. Critical constants or selected gases.From Nordstrom and Munoz (1994).

Tc(K) Pc(bar) Vc(cm3/mol) Zc

H2O 647.3 220.4 56.0 0.229

CO2 304.2 73.7 94.0 0.274

CH4 190.6 46.0 99.0 0.288

H2S 373.2 89.3 98.5 0.284

SO2 430.8 78.8 122.0 0.268

HF 461.0 65.0 69.0 0.120

HCl 324.6 83.1 81.0 0.249

NH3 405.6 112.7 72.5 0.242

H2 33.2 13.0 65.0 0.305

O2 154.6 50.4 73.4 0.288

N2

150.8 48.7 74.9 0.291


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varies insignicantly with depth in these systems and so does pressure due to the low density othe vapor. Both have values close to those corresponding to maximum steam enthalpy (234 C

and 30.1 bar absolute). Vapor-dominated systems are overlain by a cap rock, which may beliquid saturated. Vapor-dominated systems are considered to orm by the boiling down o anearlier liquid-dominated system with buoyancy driven segregation o the vapor rom the boilingliquid. The extent o vapor ormation is probably determined by the balance between the rateo fuid circulation and the rate o heat transer rom the rock to the fuid in the root zones othe system. The best known vapor-dominated elds in the world are Larderello in Italy and TheGeysers in Caliornia. Part o the Olkaria geothermal system in Kenya had prior to exploitationa steam cap o ~200 m thickness overlying a liquid-dominated system. Pressure drawdowncaused by the exploitation o the Wairakei and Svartsengi elds in New Zealand and Iceland,respectively, has led to the ormation o steam caps on top o the liquid-dominated systems(Clotworthy 2000; rmannsson 2003). In both these elds, as well as at Olkaria, the vapor capshave temperatures and pressures close to those o maximum enthalpy or saturated vapor. AtWairakei, the liquid-dominated system below the vapor cap is not much above 240 C and atSvartsengi, it is close to 240 C at least down to ~2000 m depth.

Boiling and fuid phase segregation in wells and producing aquiers

Aquiers that are penetrated by wells drilled into high-temperature, liquid-dominatedgeothermal systems are oten at P-Tconditions insucient to initiate boiling in the containedgeothermal fuid. I the pressure drop caused by discharging such wells is not sucient toinitiate boiling in the aquier itsel, the depth level o rst boiling is within the well. Under theseconditions, it is a reasonable approximation to treat the aquier and well as an isolated system,in which case boiling is adiabatic. According to thermodynamics, a system is isolated when notranser o heat or mass occurs across its boundaries. Hence, considering well and aquier asan isolated system implies that the discharge enthalpy and chemical composition o the welldischarge are the same as those o the aquier water. By mass conservation, we then have

M M M Mf t d t d v d l, , , , ( )= = + 11

whereMf,t,Md,t,Md,v, andMd,l designate the mass fow rate o aquier water into the well, mass

fow rate o the well discharge, mass fow rate o vapor rom the well and mass fow rate oliquid rom the well, respectively. The conservation o mass o dissolved components may beexpressed by

m m m X m X if t

i

d t

i

d v d v

i

d l d v, , , , , ,( ) ( )= = + 1 12

whereXd,v is the vapor mass raction o the well discharge and thereore (1 Xd,v) is the liquidmass raction. Also, mi represents the molal concentration o dissolved component i. The super-scripts have the same meaning as in Equation (11). For the specic enthalpy (h) we have

h h h X h X f t d t d v d v d l d v, , , , , ,( ) ( )= = + 1 13

The enthalpy o saturated steam is the sum o the enthalpy o steam-saturated liquid and thelatent heat o vaporization (L). Inserting h d,l +Ld or h d,v into (13) and rearranging to isolateXd,v yields

X h hL

d v

d t d l

d,

, ,

( )= 14

Values or h d,l and Ld can be obtained rom Steam Tables and a value or h d,t can either beobtained rom measurement o the discharge enthalpy or, in the case o sub-boiling aquier,rom evaluation o the aquier temperature. In a sub-boiling aquier, the enthalpy o the aquierfuid (hf,t) is simply that o liquid water (hf,l) at the aquier temperature and pressure and can beound in standard Steam Tables. Having obtained h d,t, h d,l andLd, the steam raction (Xd,v) can


16/54


be retrieved rom Equation (14). Finally, Equation (12) is used to calculate the concentration ocomponent i in the aquier water (mi

f t, ) rom analysis o liquid (mid l, ) and vapor (mi

d v, ) samples

collected at the wellhead at vapor pressure Pd.When a well intersects an aquier that is two-phase (both liquid and vapor are present),

intensive depressurization boiling will start in this aquier when the well is discharged. In casethe pressure drop produced by discharging the well is suciently large, boiling can also startin the aquier even i the aquier fuid was initially non-boiling. When intensive boiling occursin producing aquiers, vapor may not only orm by depressurization boiling. It can also ormby conductive heat transer rom the rock to the fuid: Depressurization boiling lowers the fuidtemperature, thus creating a temperature gradient between fuid and aquier rock and avoringconductive heat transer rom rock to fuid. Addition o heat to the two-phase fuid will, ocourse, not aect its temperature. It will enhance boiling, i.e., steam ormation. Liquid andvapor may segregate (separate partly) in two-phase aquiers, leading to an increase in the vaporto liquid ratio o well discharges. Phase segregation results rom the dierent fow propertieso liquid and vapor and the eects o capillary pressure and relative permeability (e.g., Horneet al. 2000; Pruess 2002; Li and Horne 2004). The mass fow rate o each phase is aected bythe relative permeability and the pressure gradient, and also its density and viscosity. Adhesiveorces between mineral grain suraces and fuid, which are the cause o capillary pressure,are stronger or liquid than or vapor. In this way, the mobility o liquid is reduced relative tothat or vapor. The eect o capillary pressure becomes stronger in rocks o small pores andractures, i.e., when permeability is low. Typical relative permeability curves are shown inFigure 9. They show the relationship between liquid saturation (volume raction o total fuidthat is liquid) and relative permeability. It can be seen rom the gure that the liquid phase isimmobile when its volume raction is below about 0.6. Relative permeability curves, such asthose shown in Figure 9, vary with temperature and interace between fuid and rock.

Due to vapor-orming processes and phase segregation in two-phase aquiers, the dischargeenthalpy o wells producing rom liquid-dominatedgeothermal reservoirs is oten higher than theenthalpy o the initial aquier fuid, and it is not uncommon that wells drilled into such systemsdischarge steam only. Wells with discharge enthalpies higher than that o steam-saturated water

at the aquier temperature have been reerred to as excess enthalpy wells. Some o the excessenthalpy may be due to the presence o vapor in the initial aquier fuid. However, the eects o

0.292 1 0.566 0

0.312 0.965 0.64 0.01

0.34 0.88 0.7 0.025

0.4 0.7 0.74 0.038

0.48 0.485 0.8 0.06

0.51 0.41 0.86 0.11

0.54 0.35 0.9 0.175

0.6 0.232 0.94 0.29

0.64 0.175 0.98 0.525

0.74 0.056 1 0.85

0.84 0.012

0.883 0

0

0.2

0.4

0.6

0.8

1

0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1

Liquid saturation

Relativepermeability

vapor

liquid

Figure 9. Typical relative permeability unctions. They show the relationship between liquid saturation(raction o liquid water by volume) and relative permeability. It can be seen rom this gure that the liquidwater phase is immobile when water saturation is below about 0.6. From Pruess (2002).


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conductive heat transer rom the aquier rock to the depressurized fuid, or phase segregation,or both, generally seem to be more important. To calculate the chemical composition o the

initial liquid and vapor in the aquier, the relative contributions o the dierent processes to theexcess discharge enthalpy and the initial vapor raction o the aquier fuid need to be evaluated.The calculation procedure, which is rather complex, is detailed in Appendix I.

In general, wet-steam wells and their producing aquiers may be dened as isolated,closed or open systems1. Five models are considered in the ollowing, three o which describeopen systems. All ve are presented in Figure 10. They dier by changes the fowing fuidundergoes with respect to enthalpy and composition between initial aquier and well, asshown by the arrows in Figure 10. The processes that change the fuid are vapor ormationby depressurization boiling, additional vapor ormation by conductive heat transer romthe rock to immobile capillary water, enhanced vaporization o the fowing liquid, also byconductive heat transer rom the rock, and fuid phase segregation leading to the retention osome o the boiled liquid in the aquier by its immobilization. Conductive heat transer andimmobilization o boiled liquid are taken to occur at a constant selected pressure, P e. This is asimplication. One would expect both conductive heat transer and immobilization o boiledwater to occur over a range o pressures. Since the pressure at which these processes operatecannot be constrained, however, the present choice is reasonable because it is the simplest one.For specic well data, one can study how sensitive the calculated fuid parameters and thecomposition o the initial aquier liquid and vapor are to the choice o P e. It is also advisableto study how sensitive the calculated parameters are to the selection o aquier temperature.

One o our models assumes that heat transerred conductively rom aquier rock to fuidenhances evaporation o the fowing liquid water (Fig. 10, Model 2). In physical terms, thismodel is logical when the capillary water has been evaporated to dryness so transer o heatrom rock to fuid involves the fowing fuid.

Figure 10 summarizes the essential eatures o the ve models, and Table 4 lists all theirparameters. Some o these are obtained rom measurement at the wellhead (Md,t, h d,t, P d, mid,l,mi

d,v), others are selected (Tf, P e) and still others are calculated (Xf,v, mif,l, mif,v, Qe,Me,v, Me,l,Mf,t).

The various models may be specied by setting one or more o the fuid parameters tozero. In Model 1 (see Fig. 10), there is no conductive heat transer to the fuid (Qe = 0), noretention o liquid by the ormation (Me,l = 0), and no additional infow o vapor (Me,v = 0). InModel 2, there is neither infow o additional vapor nor retention o liquid by the ormation(Me,v = 0 andMe,l = 0). There is, however, conductive heat fow to the fuid (Qe 0). Model 3includes no heat transer (Qe = 0) and no addition o vapor (Me,v = 0), but liquid is retained bythe ormation (Me,l 0). Model 4 involves the addition o vapor to the fowing fuid (Me,v 0)and the retention o liquid by the ormation (Me,l 0), but no conductive heat transer (Qe = 0).In Model 5, water is retained by the ormation (Me,l 0) and there is conduction o heat (Qe0), but no addition o vapor (Me,v = 0).

Glover et al. (1981) used the chemistry o well discharges in a simple way to distinguishbetween excess well enthalpy caused by phase segregation on one hand, and by conductive heattranser rom aquier rock to fuid on the other. I the concentration o a non-volatile, conservative

component like Cl

in the total discharge o a well stays about constant despite variations inexcess discharge enthalpy, the cause o the excess enthalpy is conductive heat transer rom theaquier rock to the fowing fuid (closed system, Fig. 10). I the Cl concentration stays constantin the total discharge, it will become very high in the liquid phase as the enthalpy approachesthat o saturated steam (Fig. 11A). I, on the other hand, the concentration o a conservative

1 In the geothermal industry, the term wet-steam well denotes a well that discharges a mixture o liquid andvapor.


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Depre

ssuri

zatio

n

zone

Aquif

er

We

ll

P

h

m

m

M

d

d,t

d,v

d,l

d,t

i

i

Initialaquiferconditions

T

m

m

f

f,v

f,li

i

Model 1

Depre

ssuri

zatio

n

zone

Aquif

er

We

ll


Model 3

Me,l

Depre

ssuri

zatio

n

zone

Aquif

er

We

ll


M Me,v e,l

Model 4

Depre

ssuri

zatio

n

zone

Aquif

er

We

ll


Qe

Model 2

Depress

uriza

tion

zone

Aquif

er

We

ll


Q Me e,l

Model 5

T

m

m

f

f,v

f,li

i

T

m

m

X Mf,v f,t

T

m

m

f

f,v

f,li

i

Tm

m

f

f,v

f,li

i

T

m

m

f

f,v

f,li

i

X Mf,v f,t

X Mf,v f,t

X Mf,v f,t

X Mf,v f,t

P

h

m

m

M

d

d,t

d,v

d,l

d,t

i

i

P

h

m

m

M

d

d,t

d,v

d,l

d,t

i

i

P

h

m

m

M

d

d,t

d,v

d,l

d,t

i

i

P

h

m

m

M

d

d,t

d,v

d,l

d,t

i

i

Figure 10. Schematic diagrams o the vemodels used to calculate aquier steamractions and initial aquier liquid andvapor compositions or wet-steam wells.The various parameters and symbolsare presented in Table 4 and the List o

Symbols. The equations describing themodels are given in Appendix I.


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component in either the liquid (e.g., Cl) or vapor (e.g., Ar) stays constant despite variationsin excess enthalpy, the cause is phase segregation. Constant concentration o a conservativecomponent that only occupies the liquid, collected at a particular vapor saturation pressure,implies that its concentration in the total discharge approaches zero as the discharge enthalpyapproaches that o saturated steam (Fig. 11B).

For many wet-steam wells, discharge enthalpy may not vary suciently with time to

make the method o Glover et al. (1981) applicable. In a specic welleld, one may use manywells with a range o discharge enthalpies to determine whether conductive heat transer orphase segregation dominates the excess enthalpy. I components like Cl vary much acrossthe welleld, they may not be a good choice, but i aquier temperatures are about constant,SiO2 is useul because its concentration is determined almost solely by temperature throughits equilibrium with quartz. As an example, data rom Olkaria in Kenya show how SiO2concentrations in the total discharge o wells vary with discharge enthalpy (Fig. 12). They

Table 4. Measured, selected and calculated fuid and composition parameters by the vemodels considered in this contribution. Explanation o parameter symbols are given in

Figure 10 and under List o Symbols.

Model 1 2 3 4 5

Measuredparameters

Md l,

hd t,

mid l,

Md v,

Pd

mid v,

Md l,

hd t,

mid l,

Md v,

Pd

mid v,

Md l,

hd t,

mid l,

Md v,

Pd

mid v,

Md l,

hd t,

mid l,

Md v,

Pd

mid v,

Md l,

hd t,

mid l,

Md v,

Pd

mid v,

Selectedparameters T

fT

f T

eT

fT

fT

eT

fT

e

Calculatedparameters

Xf v,

mif l, mi

f v,

Xf v,

mif l,

Qe

mif v,

Xf v,

Mf t,

mif l,

Me l,

Te

mif v,

Xf v,

Mf t,

mif l,

Me l,

Me v,

mif v,

Xf v,

Mf t,

mif l,

Me l,

Qe

mif v,

0

5000

10000

15000

20000

1000 1500 2000 2500

Cl(ppm)

Discharge enthalpy (kJ/kg)

Cl in water collected at

atmospheric pressure

Cl in total

discharge

0

200

400

600

800

1000

1200

1000 1500 2000 2500

Cl(ppm)

Discharge enthalpy (kJ/kg)

Cl in total

discharge

Cl in water coll ected at

atmospheric pressure

A B

Figure 11. Relation between Cl and discharge enthalpy. A: Excess enthalpy is due to conductive heattranser rom aquier rock to fuid fowing into well (closed system). B: Excess enthalpy is caused by phasesegregation in the producing aquier (open system).


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show that SiO2 concentrations in the total discharge approach zero when the discharge enthalpyapproaches that o saturated steam, indicating that the excess enthalpy is mainly caused byphase segregation.

INITIAL AND EQUILIBRIUM VAPOR FRACTIONS

The vapor raction that may be present in an aquier fuid, prior to induction o extensiveboiling by discharging a well, is called the initial vapor raction. Geochemical methods havebeen used to evaluate this vapor raction rom the gas content o well discharges, assumingeither specic gas-gas or mineral-gas equilibria. When estimated in this way, the initial vaporraction has appropriately been termed equilibrium vapor raction.

Giggenbach (1980) estimated equilibrium vapor ractions in the reservoir o threeNew Zealand geothermal systems based on the assumption o equilibrium or the ollowingreactions:

CH4 + 2H2O = CO2 + 4H2 (15)

and

2NH3 = N2 + 3H2 (16)

The model used by Giggenbach (1980) allows the presence and loss o equilibrium vapor. Hismodel may thereore be considered as a special type o an open model, i.e., one which permitsloss or gain o gases rom the initial aquier fuid and a corresponding loss or gain o enthalpy.The retrieved equilibrium vapor raction values (Xf,v) ranged rom 5% to +1.2% by weight(~68% and ~+33% by volume), negative values indicating loss o equilibrium vapor rom thefuid entering the wells. By assuming Xf,v to be zero, Giggenbach (1980) calculated gas-gasequilibrium temperatures or the reactions given above. Gas equilibrium temperatures basedon the two reactions considered compare very well. On the other hand, correlation betweenquartz equilibrium and gas geothermometer temperatures is poor. The model adopted byGiggenbach (1980) did not consider how the excess well enthalpy had developed. This is notimportant or most o the wells considered in his study, because their discharge enthalpy isclose to liquid enthalpy. Two o the wells included in Giggenbachs study, however, dischargeddry vapor. The calculated equilibrium vapor raction or these wells is signicantly aected

0

100

200

300

400

500

0 1000 2000 3000

Total well discharge enthalpy (kJ/kg)

SiO2

intotalwelldischarge(mg/kg)

Figure 12. Variation in the silica contentin total well discharges o productionwells in the Olkaria East production eldKenya. The enthalpy value approachesthat o saturated steam when the silicacontent approaches zero. The observedcorrelation indicates that excess enthalpyo well discharges is mainly caused byphase segregation in producing aquiers.


21/54


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to represent the concentrations o CO2, H2S and H2 in the initial aquier water. The equilibriumvapor ractions are in the range o 0.38 to +0.25 wt%, the average being 0.03. Their results

indicate that a relatively large part o the vapor discharged rom wells is generated romconductive transer o heat rom the aquier rock to capillary liquid. The results o Arnrsson etal. (1990), described above, are sensitive to the equations selected to describe the equilibriumconcentrations o CO2, H2S and H2 in the initial aquier liquid. According to Karingithi (pers.comm.), equilibrium vapor ractions in the aquier o selected Olkaria wells are below 0.1wt%. This estimate should be considered more reliable than earlier estimates o Arnrssonet al. (1990), because the mineral-gas equilibrium constants used by Karingithi are based onanalyses o the chemical composition o the minerals in the mineral-gas buers and on the mostrecent thermodynamic data bases (Holland and Powell 1998; Robie and Hemingway 1995).Gudmundsson and Arnrsson (2002) estimated the equilibrium vapor raction or some Krafawells in Iceland by the method o Arnrsson et al. (1990). They obtained values in the range o0.07 to +2.22 wt% (corresponding to 0 to 47 % by volume).

The mole ractions o CO2 in the vapor discharged rom wells in the Larderello and TheGeysers vapor-dominated elds correspond to CO2 partial pressures o about 0.7 and 0.1 bar,respectively, according to data presented by Ellis and Mahon (1977). These values correspondwell with equilibrium between fuid and the mineral buer clinozoisite + prehnite + calcite +quartz (see Eqn. 1 in Table 7) with an assumed clinozoisite activity o ~0.1. This indicates thatthe vapor o these vapor-dominated systems is close to equilibrium with hydrothermal mineralstypically ound in liquid-dominated systems with temperatures o ~240 C, i.e., the vapor isequilibrium vapor. The intergranular liquid (see Fig. 2) has apparently closely approachedequilibrium with the mineral buer in question, and the vapor has rapidly equilibrated withthis liquid.

GAS CHEMISTRY

Gas chemistry o geothermal fuids is intimately related to boiling o geothermal fuids.Specic gas ratios or gas-gas reactions have been used to model aquier fuid compositions

rom data on the gas content o vapor samples collected at the wellhead. Gas concentrationsand gas ratios have also been used as geothermometers using both well and umarole data (e.g.,Giggenbach 1981; DAmore and Celati 1983; DAmore and Truesdell 1985; Arnrsson et al.1990). Finally, as discussed in the previous section, gas chemistry has been used to estimate theinitial vapor raction in aquiers producing into wet-steam wells.

Various conventions have been used to present gas compositions o geothermal fuiddischarges. They include expressing gas concentration in the vapor phase or as mole raction othe gas phase, usually as mole percent (volume percent). Various concentration units have alsobeen used (see Henley et al. 1984). These include millimoles/kg o H2O and millimoles/100 moleso H2O. A logical selection o concentration units is that used in chemical thermodynamics, i.e.,moles (or millimoles) per kg o solvent H2O, because today gas data interpretation requentlymakes use o the principles o chemical thermodynamics. When a separate gas phase is collected,the only choice is to express analytical results in terms o mole percent (volume percent). Thisis, e.g., the case when samples are collected o gas bubbling through hot spring water or mud

pools. It is, however, advantageous to collect samples rom umaroles and well discharges insuch a manner that the concentration in the vapor is recorded, because this provides importantadditional inormation or data interpretation.

The gas composition o geothermal fuids is highly variable. In volcanic geothermal systemsCO2 is the major gas constituent (Table 5). It may exceed 10% o the volume o the vapor phasebut typically it is


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are very variable. They seem to be highest in fuids that have reacted with marine sediments. Insome geothermal systems, N2/Ar ratios are close to those o air-saturated water (e.g., samplesno. 3 and 8 in Table 5), but in others they are much higher (e.g., sample no. 9 in Table 5).Elevated N

2/Ar ratios relative to those o atmospheric air may be a refection o organic source

(see, e.g., Mariner et al. 2003; Snyder et al. 2003). In areas o andesitic volcanism on convergingplate boundaries, this organic source may ultimately be marine sediments containing organicmatter on top o the downward moving lithosphere plate. The andesitic magma may assimilatethese sediments, gaining N2 and other chemical components in the process.

Except or He, noble gas concentrations in geothermal fuids typically are controlled bytheir concentrations in the source fuid, i.e., air-saturated water (e.g., Mazor and Fournier 1973;Mazor and Truesdell 1984). The concentrations o He, and sometimes also those o Ar, may beelevated due to radiogenic sources. Elevated 3He/4He ratios (5 to over 30 times atmospheric)are typical or volcanic geothermal systems due to the supply o3He rom mantle sources. Thehighest ratios are ound in hot-spot areas such as Hawaii and Iceland (e.g., Welhan et al. 1988;Poreda et al. 1992; Hulston and Lupton 1996).

The gas content o geothermal discharges (umaroles and wells) has been used to obtaininormation on the source o the fuid and its temperature. Arnrsson (2000) give a summary

o gas geothermometry. Table 6 provides geothermometer equations or the reactions inTable 7. Gas geothermometers may be based on assumptions o specic gas-gas or mineral-gas equilibria or a distribution o isotope ratios between gaseous species. Gas geothermometersinclude both gas concentrations and gas ratios or a combination o ratios. Early calibrationso gas geothermometers were empirical, i.e., based on drillhole data. Empirical calibrationinvolves correlating specic gas concentrations, gas ratios or a combination o ratios in welldischarges with the aquier temperature o the wells (DAmore and Panichi 1980; Arnrsson and

Table 5. Gas concentrations (mmoles/kg) in selected geothermal wells.

Locationa hf Pd Tqtz CO2 H2S H2 CH4 NH3 N2 Ar1 Asal, Djibouti, 6 36.1 0.07 0.06 0.033 0.74 0.011

2b Broadlands, NZ, 25 1420 16.2 284 291.1 0.74 2.23 8.02 2.23 3.63

3 Cerro Prieto, Mex, 5 1284 4.1 289 275.2 15.7 10.1 13.60 5.14 1.74 0.042

4 Kawerau, NZ, 7 1030 7.3 258 47.3 1.93 0.03 0.204 1.68 0.38

5 Krafa, Iceland, 20 2306 9.8 280 887.0 40.5 24.7 0.320 2.78 0.049

6 c Larderello, Italy, 1 726.9 13.3 18.0 8.600 4.10

7 Mahanagdong, Phil, 9 1287 3.8 278 222.0 13.5 0.05 0.360 0.03 7.03

8b Momotombo, Nic., 4 1257 0.0 262 297.7 5.96 0.40 0.449 14.91 0.015

9 Nmajall, Icel, 11 1850 25.3 260 142.0 51.4 89.5 0.782 2.47 0.050

10 Olkaria, Kenya, 2 1839 4.8 251 101.0 5.61 2.94 0.722 1.94

11 Olkaria, Kenya, 301 1653 1.5 262 4260 3.55 1.05 0.833 11.27

12 Reykjanes, Icel, 15 1256 25.0 260 389.7 10.3 0.78 0.067 7.88 0.122

13 Tongonan, Phil, 510 1376 11.1 251 139.3 10.7 1.08 0.96 0.75 5.90 0.06714 Zunil, Guatemala, 3 1423 7.3 265 130.8 5.63 0.23 0.017 0.25 1.36 0.006

15 Wairakei, NZ, 72 1295 9.0 246 13.8 0.54 0.04 0.152 0.68 0.22

hf: discharge enthalpy, kJ/kg; P d: sampling pressure, bar-g; Tqtz: quartz equilibrium temperature, C.aThe numbers ater the local name in column 2 reer to well number. bNot the same well rom which data are presentedin Table 1 but a well rom the same area. cDry steam well.

Source o data: 1 (1): DAmore et al. (1998); 2 (8), 4, 15 (7): Giggenbach (1980); 3: Nehring and DAmore (1985); 5(11), 9 (12): Gudmundsson and Arnrsson (2002); 6: DAmore and Truesdell, 1984; 7 (6), 13 (3): Angcoy (pers. comm.);8 (4): Arnrsson (1997); 10 (9), 11 (10): Karingithi (pers. comm); 12 (2): Giroud (pers. comm.); 14 (5): Arnrsson(1995b). The numbers in parentheses above reer to the numbers in Table 1 or water sample analysis.


24/54


Gunnlaugsson 1985). Calibration has also been based on thermodynamic data or either specicgas-gas or mineral-gas reactions (Giggenbach 1980; Nehring and DAmore 1984; Arnrsson etal. 1998; Karingithi pers. comm.). Calibrations based on thermodynamic data always rest on anassumption o specic chemical equilibria. This assumption needs to be veried by comparinggas geothermometer results with solute or isotope geothermometers and especially with welltemperature data. In the upfow zones o geothermal systems, the vapor may condense partly,either by conductive heat loss or by fowing through water in perched aquiers. Condensationincreases the gas concentrations in vapor. To cancel the eect o condensation by conductive

Table 6. Temperature equations or gas (steam) geothermometers basedon the reactions in Table 7. Arnrsson et al. (2000) provide detailed

summary on gas geothermometry equations.

The temperature equations below correspond to the reactions given in Table 7. Theyare valid in the range 150-350 C. All minerals appearing in reactions 1 to 10 in Table7 are taken to be pure (their activity equals 1), except or epidote, prehnite and garnetthat orm solid solutions. To apply these geothermometers requires knowledge o theacticity o the clinozoisite, Al-prehnite and grossular end members in the respective solidsolutions rom their analysed compositions.

Reactiona

1 Ta apre czo( )

log( ) . . log( ) log( )

.C

CO2=+ + 3 28 1 5

0 0097

2 Ta agro czo

( )

log( ) . . log( ) . log( )

.C

CO2=

+ + 5 072 0 6 0 4

0 01425

3 Ta aepi pre( )

log( ) . ( / ) log( ) ( / ) log( )

.C

H S2=+ + 6 853 2 3 2 3

0 013343

4 Ta awol epi( )

log( ) . ( / ) log( ) ( / ) log( ) ( / C

H S2=+ + + 6 722 2 3 2 3 2 33

0 01394

) log( )

.

agro

5 Ta aepi gro( )

log( ) . log( ) log( )

.C

H S2=+ + 6 571 2 2

0 01664

6 T( )log( ) .

.C

H S2=+ 5 537

0 01197

7 T a aepi pre( ) log( ) . ( / ) log( ) ( / ) log( ).C H2= + + 4 686 2 3 2 30 0079962

8 Ta aepi gro( )

log( ) . ( / ) log( ) ( / ) log( )

.C

H2=+ + 4 675 2 3 2 3

0 0076623

9 Ta aepi gro( )

log( ) . log( ) log( )

.C

H2=+ + 7 242 6 6

0 01939

10 T( CH2 )

log( ) .

.=

+ 4 3310 005719

aSee Table 7.


25/54


heat loss, Giggenbach (1988) calculated the ratios o the reactive gas concentrations (CO2 andH2) to the inert gas Ar. Arnrsson (1987) proposed a correction procedure that also involved

partial vapor condensation, both by conductive heat loss and mixing with ground water. In thecase o mixing, degassing o the liquid was taken into account.

Relatively ew attempts have been made to veriy or disprove whether specic gasequilibria are closely approached in the aquiers o wet-steam wells. It appears, that equilibriumis requently closely approached between mineral buers and CO2, H2S and H2 in the aquierso wet-steam well discharges (e.g., Arnrsson and Gunnlaugsson 1985; Gudmundsson andArnrsson 2002; Karingithi pers. comm.). How closely equilibrium is approached, dependson the gas fux rom the magma heat source. Thus, in the western part o the Olkaria eldin Kenya, CO2 concentrations are very high in well discharges and apparently not controlledby equilibrium with a mineral buer. This has been attributed to a high fux o this gas romthe magmatic source. In the eastern hal o Olkaria, equilibrium between CO 2 and a mineralbuer seems, however, to be closely approached (Fig. 13). One important eature can beobserved in Figure 13, which is that dierent mineral buers give very similar aqueous CO 2,H2S and H2 concentrations at equilibrium when the end-member activities in the solid solutionminerals are properly chosen. The same mineral buer is likely to constrain H 2S and H2 aquierliquid concentrations at equilibrium. The studies o Arnrsson and Gunnlaugsson (1985)and Gudmundsson and Arnrsson (2005) indicate that CO2 in


26/54


Table7.Equilibriumconstants(asdissolvedgasandpartialp

ressure)ormineralgasreactionsthatcan

potentiallycontrolaqueousCO2,H2SandH

2

concentrationsingeothermalsystems.Theyarevalidintherange0350CatPsat.Unitactivitywas

selectedorallmineralsandliquidwater.

Gas

Reaction

logK

(T)

1

CO2

czo

cal

qtz

pre

l

aq

+

+

+

=

+

3 2

3 2

2

H

O

CO

2

,

1.83122843/T21344.4/T+0.00829T0.455logT

2

CO2

2 5

3 5

3 5

1 5

2

czo

cal

qtz

gro

l

aq

+

+

=

+

+

H

O

CO

2

,

2.32758215/T21829.2/T+0.01059T0.561logT

3

H2S

1 3

1 3

2 3

2 3

2

2 3

2

H

O

H

S

pyr

pyrr

pre

epi

l

aq

+

+

+

=

+

2.052215218/T21658.7/T

+0.00856T0.509logT

4

H2S

2 3

1 3

1 3

2 3

4 3

2

2 3

2 3

2

H

O

H

S

gro

pyr

pyrr

qtz

epi

wol

l

aq

+

+

+

+

=

+

+

1.862253145/T21595.2/T

+0.00780T0.452logT

5

H2S

2

2

2

2

2

gro

pyr

mag

qtz

epi

wol

l

aq

+

+

+

+

=

+

+

1 4

1 2

2

2

H

O

H

S

1.549383405/T21774.0/T

+0.00820T0.303logT

6

H2S

1 4

1 2

2

1 4

2

H

O

H

S

pyr

pyrr

mag

l

aq

+

+

=

+

2.020188233/T21504.4/T

+0.00760T0.528logT

7

H2

4 3

2 3

2 3

2

2 3

2 3

H

O

H

pyrr

pre

epi

pyr

l

aq

+

+

=

+

+

2,

1.3581420.4/T+6.777104T+5.611106T20.391logT

8

H2

2 3

4 3

2 3

4 3

2

2 3

2 3

2 3

H

O

H

gro

pyrr

qtz

epi

wol

pyr

l

aq

+

+

+

=

+

+

+

2,

1.2411519.9/T+8.818104T+4.693106T20.336logT

9

H2

6

2

6

4

6

6

2

gro

mag

qtz

epi

wol

l

aq

+

+

+

=

+

+

H

O

H

2

,

1.5705346.3/T+0.00880T

6.479106T2+1.113logT

10

H2

2 3

2

3 4

1 4

H

O

H

pyrr

pyr

mag

l

aq

+

=

+

+

2,

1.4361131.3/T1.866104T+5.377106T20.454logT


27/54


Gas

Reaction

logPgas(bars)

1

CO2

czo

cal

qtz

pre

l

aq

+

+

+

=

+

3 2

2

3 2

H

O

CO

2,

57.78122843/T24792.99/T

+0.00829T+0.6864106T219.302logT

2

CO2

2 5

3 5

3 5

1 5

2

H

O

CO

czo

cal

qtz

gro

l

aq

+

+

=

+

+

2,

57.28558215/T25277.79/T

+0.01059T+0.6864106T219.408logT

3

H2S

1 3

1 3

2 3

2 3

2

2 3

2

H

O

H

S

pyr

pyrr

pre

epi

l

aq

+

+

+

=

+

66.723215218/T25331.78/T+0.00856T+4.07153106T223.070logT

4

H2S

2 3

1 3

1 3

2 3

4 3

2

2 3

2 3

2

H

O

H

S

gro

pyr

pyrr

qtz

epi

wol

l

aq

+

+

+

+

=

+

+

66.913253145/T25268.28/

T+0.00780T+4.07153106T223.013logT

5

H2S

2

2

2

2

2

gro

pyr

mag

qtz

epi

wol

l

aq

+

+

+

+

=

+

+

1 4

1 2

2

2

H

O

H

S

67.226383405/T25447.08/

T+0.00820T+4.07153106T222.864logT

6

H2S

1 4

1 2

2

1 4

2

H

O

H

S

pyr

pyrr

mag

l

aq

+

+

=

+

66.755188233/T25177.48/

T+0.00760T+4.07153106T223.089logT

7

H2

4 3

2 3

2 3

2

2 3

2 3

H

O

H

pyrr

pre

epi

pyr

l

aq

+

+

=

+

+

2,

23.9022775.68/T+6.77710

4T+1.49953106T27.357logT

8

H2

2 3

4 3

2 3

4 3

2

2 3

2 3

2 3

H

O

H

gro

pyrr

qtz

epi

wol

pyr

l

aq

+

+

+

=

+

+

+

2,

24.0192875.18/T+8.81810

4T+0.58153106T27.302logT

9

H2

6

2

6

4

6

6

2

gro

mag

qtz

epi

wol

l

aq

+

+

+

=

+

+

H

O

H

2

,

26.8306701.58/T+0.00880T

10.59047106T25.853logT

10

H2

3 2

2

3 4

1 4

H

O

H

pyrr

pyr

mag

l

aq

+

=

+

+

2,

23.8242486.28/T1.86610

4T+1.26553106T27.420logT

ThethermodynamicpropertiesoCO2,aq,H2SaqandH2,aqareromFernandezPrinietal.(2003).MineraldataareromHollandandPowell(1998)exceptorpyriteandpyrrhotitethatare

rom

RobieandHemingway(1995).DataonH4SiO4

0andquartzsolubilityarerom

GunnarssonandArnrsson(2000),thoseonH2Ol,Ca

+2,Fe+2andOHromSUPCRT92program(Johnson

etal.

1992)usingtheslop98.datdataset.ThedataonFe(OH)4andAl(OH)4areromDiakonovetal.(1999)andPokrovskiiandHelgeson

(1995),respectively.Thefuoritesolubilityequationis

rom

Arnrssonetal.(1982).ItisbasedonNordstromandJenne(1977).Abbrevatio

nsormineralphasesare:cal:calcite;czo:clinozoisit

e;epi:epidote;fu:fuorite;gro:grossular;mag:magn

etite;

pre:prehnite;pyr:pyrite;pyrr:pyrrhotite;qtz:quartz;wol:wollastonite.


28/54


{{ {{{{{

{{

{{{{

{{{{

{ {

{

{

{

{{

{

{{

{{{

-3

-2

-1

0

1

2

150 200 250 300

logP(CO2

)

Temperature (C)

1

2

{{ {{{

{{

{{

{{

{{{{{

{{ {

{

{

{

{

{

{

{{{

{

{

-4.5

-3.5

-2.5

-1.5

-0.5

150 200 250 300

logP(H2

S)

Temperature (C)

1

2

34

{{

{

{{{{ {{{{

{{

{{{{{ {

{

{{

{

{{{

{

{

{

-5

-4

-3

-2

-1

0

150 200 250 300

logP(H2

)

Temperature (C)

1

23

4

1 clinozoisite+calcite+quartz+prehnite

2 clinozoisite+calcite+quartz+grossular

1 epidote+prehnite+pyrite+pyrrhotite

2 magnetite+pyrite+pyrrhotite

3 epidote+grossular+pyrite+pyrrhotite+quartz+wollastonite

4 epidote+grossular+magnetite+pyrite+quartz+wollastonite

Figure 13. Partial pressures o CO2, H2S and H2 in the initial aquier fuid producing into wells drilled intothe Olkaria geothermal eld in Kenya (circles). The curves represent equilibrium constants or selectedmineral buer-gas reactions as shown. In deriving these curves the average analyzed compositions o epidote,

prehnite and garnet in the Olkaria system were taken into consideration. Other minerals (calcite, quartz,pyrite, pyrrhotite) in the respective buers were taken to be pure. The Olkaria data indicate that equilibriumis closely approached in most aquiers between the aquier water and the mineral buers epidote + prehnite +calcite + quartz (or CO2) and epidote + prehnite + pyrite + pyrrhotie (in the case o H2S and H2). However,some well discharges (in the west part o the Olkaria eld) have CO 2 concentrations higher than thosecorresponding to equilibrium. It is thought that these high CO2 concentrations are determined by a high fuxo this gas rom the magmatic heat source rather than by local mineral buer equilibrium. Based on datarom Karingithi (pers. comm.).


29/54


hanced gas fux into a non-boiling geothermal system would shit the depth level o rst boilingto greater depths. In both instances the enhanced gas fux will cause some cooling o the fuid.

BOILING AND CHANGES IN MINERAL SATURATION

Changes in fuid composition during boiling and degassing

When geothermal fuids boil extensively, two major changes occur. One involves increaseddissolved solids content o the boiled liquid due to vapor ormation. The other major changeis an increase in pH due to transer o CO2 and H2S (acid gases) rom the liquid into the

0

10

20

30

200 250 300 350

PCO

2

(bar)

Temperature (C)

1

2

0

0.5

1

1.5

2

200 250 300 350

PH

2S(bar)

Temperature (C)

0

24

6

8

10

12

200 250 300 350

PH

2(

bar)

Temperature (C)

6

4

3

5

78

9

10

Figure 14. Relationship between temperatures and gas pressures as determined by local equilibria withvarious mineral buers. The numbers correspond to the mineral buer-gas reactions shown in Table 7. Iequilibrium is attained, only CO2 and H2 will exert a suciently high partial pressure above ~300 C toaect the boiling point o water signicantly.


30/54


vapor. These changes and the cooling o the fuid by depressurization boiling cause complicatedchanges in individual aqueous species activities, as values o association constants or complexes

and ion hydrolysis constants are temperature dependent. The pH also plays a major role.The transer o dissolved gases rom the initial liquid to the vapor is determined by their

solubility in liquid and how closely equilibrium distribution is approached between the liquidand vapor. The solubility constant (Ks) o gas species s in aqueous solution is dened byEquation (4) above. The distribution coecient (Ds) or species s is dened as

Dx

xs

s

v

s

l= (18)

Here,xsv andxs

l designate the mole raction o the species in vapor and liquid, respectively. Themole raction is given by

xN

N Ns

s

i

i

=+ H O2

19( )

whereNsstands or the number o moles o volatile species s, and the sum is over all suchspecies except H2O. I we consider a mass o fuid containing 1 kg o H2O, thenNH2O = 55.51and

xm

ms

s

i

i

=+ 55 51

20.

( )

where ms denotes the molal concentration o gas s. For dilute fuids, mii


31/54


vapor samples. This is, however, neither possible downhole nor in upfow zones below boilinghot springs. As the extent o degassing during boiling aects the mineral saturation state o

the liquid, it is o interest to assess this eect. This is possible with the aid o the WATCHchemical speciation program (Arnrsson et al. 1982). A actor () is used to represent theextent o degassing o the boiled liquid relative to equilibrium degassing. In the program thisinvolves dividing Ks by the selected value (between 0.01 and 1). A value o 1 correspondsto equilibrium gas distribution between the liquid and vapor. A small value indicatesnegligible degassing. An equation analogous to Equation (12) permits calculation o the liquidconcentration o gas species s at any selected vapor pressure. We have

m m X m X st

s

v v

s

l v= + (26)( )1

Combination o (22), (25) and (26) and insertion o yields

m m XP K

s

l

s

t l v

tot s

=

+

(27),.55 51

1 1

1

For Equation (27) to be useul,Xv must be known. For wet-steam wells, a value orXv canbe obtained rom the measured discharge enthalpy (h d) applying Equation (A4). For boilinghot spring fuids, Xv can be obtained rom selection o an aquier temperature by taking theenthalpy o the parent aquier fuid (hf,t) to be equal to that o vapor saturated liquid (hf,l) atthe aquier temperature.

Mineral deposition with special reerence to calcite

In upfow zones o geothermal systems and in wells where extensive boiling occurs,mineral deposition rom the boiling fuid largely occurs in response to its cooling and degassing.Many minerals, such as amorphous silica and metallic suldes, have prograde solubility (thesolubility increases with increasing temperature). Others, including anhydrite, calcite andaragonite, have retrograde solubility with respect to temperature. Cooling causes geothermalwaters to become over-saturated with minerals with prograde solubility but under-saturatedwith those having retrograde solubility. Some minerals, such as calcite, aragonite and metallic

suldes, have pH-dependent solubility. Degassing tends to produce over-saturated water withrespect to minerals whose solubility decreases with increasing pH.

The quantity o minerals precipitated rom solution is not only determined by the degreeo over-saturation but also by the fuid composition and the kinetics o the precipitationreaction. Quartz, adularia and albite, e.g., are abundant in the boiling zone o high-temperaturegeothermal systems, but do not orm in wells due to slow kinetics. In dilute waters, the extento sulde mineral deposition is essentially limited by low metal concentrations. In brines, bycontrast, sulde mineral deposition is extensive due to the abundance in solution o cationsthat orm sulde minerals. Troublesome scale in wells is characterized by phases that readilyprecipitate rom solution, such as amorphous silica and salts including calcite, aragonite andanhydrite. Metallic suldes also precipitate readily rom solution.

The solubility o some minerals increases with increasing temperature but decreases withincreasing pH; an example being calcite. The combined eects o both processes, togetherwith the rate o the precipitation reaction, determine whether or not the minerals with whichthe water becomes over-saturated precipitate rom solution. Un-boiled geothermal liquidsare typically close to being calcite-saturated (e.g., Arnrsson 1989). Extensive degassing byboiling tends to cause an initially calcite-saturated water to become over-saturated. The coolinghas the opposite eect due to the retrograde solubility o calcite with respect to temperature.The extent o degassing and cooling determines whether boiling causes an initially calcite-saturated water to become over or under-saturated. Figure 15 shows how the degree o calcitesaturation varies during adiabatic boiling or seven selected wells, assuming maximum


32/54


(equilibrium) degassing or all analyzed gases (CO2, H2S, H2, CH4, N2). The calculated calcitesaturation in the aquier or each well is shown by dots. The departure rom saturation isvery small or the majority o the wells and is regarded to be within the limit o error when

all uncertainties behind the calculations are considered, such as analytical imprecision andselection o aquier temperature. A similar overall pattern in the variation o the saturationindex (SI) with temperature is observed or all the wells. A sharp initial increase in SI is seen,ollowed by a decline. The initial increase in SI refects an increase in pH due to CO2 and H2Sdegassing. The pH increase causes a strong increase in the activity o the CO32 species. Atmaximum SI values, the water has been largely degassed, and the subsequent decline in SI iscaused by increased calcite solubility with decreasing temperature. The only exception to thegeneral pattern just described is sample 5 in Figure 15 (Nisyros, well 2, Greece). This welldischarges a highly saline liquid (~3 times seawater salinity). The slight increase in the SI orthis well with decreasing temperature ater the initial rise in SI is largely due to dissociation oion pairs with alling temperature, mostly CaHCO3+.

Figure 16 shows how the calcite saturation index varies with the extent o degassingduring adiabatic boiling o fuid discharged rom a well in the Amatitlan eld in Guatemala.I degassing is less than 30% o maximum ( < 0.3), boiling does not produce a calcite over-

saturated solution as it does when is > 0.3.

MODELING OF AQUIFER FLUID COMPOSITIONS

This section describes modeling o fuid compositions in aquiers that eed boiling hotsprings and wells drilled into high-temperature liquid-dominated geothermal systems (wet-steam wells). It is based on analytical data on liquid and vapor samples collected at the surace.

v

-1

-0.5

0

0.5

1

100 150 200 250 300 350

SaturationIndex

Temperature C

1 2

3

4

5

6

7v

v

v

v

v

v

Figure 15. Variation in calcite saturation temperature during adiabatic boiling o fuid discharged romselected wet-steam wells. Equilibrium distribution was assumed or all gases between the liquid water andvapor. Dots indicate aquier conditions. 1: Hveragerdi, well 2, Iceland; 2: Hveragerdi, well 7, Iceland; 3:Wairakei, well 24, New Zealand; 4: Amatitlan, well 1, Guatemala, 5: Asal, well 1, Djibouti; 6: Krafa, well6, Iceland; Palinpinon, well 7, Philippines.


33/54


The modeling assumes that neither mineral precipitation nor dissolution occur betweenaquier and surace. Basically, it reconstructs the fuid composition below (beyond) the zoneo extensive boiling. Modeling o this kind is essential or studying mineral-fuid equilibriain the aquier. The model used or hot springs is relatively simple, but it is more involved orthe wet-steam wells due to the rather complicated processes in the depressurization zone thatorms around discharging wells (see Fig. 10). For hot springs, the boiling between aquier andsurace is taken to be adiabatic.

Boiling hot springs

Boiling, mixing and conductive cooling in upfow zones o geothermal systems arevariable and predominantly complex processes that may involve mineral dissolution andprecipitation reactions. At the present state o knowledge, satisactory general boiling-mixing-reaction models cannot be developed to determine initial aquier fuid compositions rom dataon hot spring discharges. In the absence o mixing, and when the temperature o the eedingaquier is not much above 200 C (or 200 C water, boiling starts at ~170 m depth), aquierfuid compositions may be reconstructed reasonably accurately below the zone o boiling,at least when fow rates are suciently high to make conductive heat loss o the rising fuidinsignicant. Under these conditions, it is logical to assume boiling to be adiabatic.

Many studies have indicated that mineral-solution equilibrium is closely approached orall major components o geothermal fuids except Cl (e.g., Giggenbach 1981; Arnrsson etal. 1983). This greatly constrains the relative activities o aqueous species, and thereore alsoindividual component concentrations, and helps evaluate the validity o the modeling results.For example, cation/proton activity ratios in equilibrated liquids are constant at any specictemperature or a system o specic mineralogy (Fig. 17).

The WATCH chemical speciation program (Arnrsson et al. 1982) version 2.1 (Bjarnason

-1

-0.5

0

0.5

100 150 200 250

SaturationIndex

Temperature (C)

0.05

0.1

0.2

0.3

0.5

1.0

v

Figure 16. Variation in calcite saturation with temperature during adiabatic boiling o fuid dischargedrom well 1 at Amatitlan, Guatemala. Degassing was assumed to be variable. The numbers by each curverepresent the selected value o the degassing coecient (). A value o 0.05 or corresponds to degassingthat is 5% o maximum, i.e., equilibrium, degassing and so on.


34/54


1994) permits calculation o aquier water composition rom data on the chemical compositiono boiling hot-spring waters. Essentially, calculations involve correcting the dissolved solidscontent o the hot spring water or vapor loss and adding back into the water those gases (CO2and H2S) that were lost with the vapor. To run the program, an aquier temperature must beselected. The choices include several geothermometer temperatures, including the temperatureo last equilibrium with quartz, and an arbitrarily selected temperature value. A value or thedegassing coecient () needs to be chosen as well. One way o doing this, is to study how relates to calcite saturation. It is logical to select a value that corresponds to calcite saturationin the aquier at the selected aquier temperature because, as already pointed out, geothermal

reservoir waters are generally close to being calcite-saturated, at least when temperatures are>100 C. The concentrations o gases other than CO2 and H2S in the un-boiled aquier fuid canbe approximated, i data are available on the relative abundance o gases in vapor samples.

To demonstrate the calculation o aquier fuid composition rom analytical data on boilinghot spring water, we have selected one such spring rom the Geysir geothermal eld in Iceland(Table 8). First the WATCH program was run or dierent values taking the temperature olast equilibrium with quartz (204 C) to represent the temperature o the un-boiled fuid below

{ {

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{

{{{

3

4

5

6

7

8

1.5 2 2.5 3 3.5

NaH

1000/TK

{ {{

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Fluid-Fluid Interactions in Geothermal Systems

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Transcript of Fluid-Fluid Interactions in Geothermal Systems