Several factors affect hydrothermal alteration and include temperature changes of boiling and cooling fluid, rock permeability, pressure, fluid composition, duration of hydrothermal activity, mineral composition of the primary rocks and chemical changes due to mixing of fluids (Pirajno, 2009). Numerous hydrothermal minerals are deposited within fissures, on wall linings of drilled wells and steam lines that in turn impact reservoir productivity. Throughout any geothermal energy producing wells, changes in temperature, pressure, pH and mineral saturation distract the chemical equilibria of the working fluid and generally lead to scale formation or corrosion. There are three main areas of secondary mineral deposition (Opondo, 2005): deposition from a single-phase fluid (injection pipelines), deposition from flashing fluid (wells, separators, two-phase pipelines) and deposition by steam carryover (separators, steam lines and turbines). When geothermal fluid ascends to the surface, hydrostatic pressure decreases. Eventually, the pressure drops toward a level that permits the dissolved gases and steam to separate from the liquid phase. This physical process of phase separation is defined as “boiling” (Nicholson, 1993). Additionally, other notable chemical processes that influence the aqueous composition include conductive cooling, dissolution and deposition (precipitation). Adiabatic flashing causes CO₂ stripping and a pH increase, which may lead to calcite scaling. The most common scales consist of calcium carbonate and amorphous silica, but scales of various mineral oxides and sulphides, mostly pyrite, also occur widely (Thomas & Gudmundsson, 1989). The precipitation of calcite from solution is largely due to the effect of steam loss and degassing of CO₂. This relationship is evident in the chemical Equation (1).

Ca ( aq ) 2 + + 2HCO 3 ( aq ) − ⇄ H 2 O ( l ) + CaCO 3 ( s ) + CO 2 ( aq ) (1)

with the head loss provided by the water flow, CO₂ changes to the gaseous state (CO_2(aq) → CO_2(g)), which disfavors the CaCO₃ dissolution and implies its mineralisation (De Paula et al., 2022). Solubility of silica minerals decreases with a decrease of temperature as opposed to carbonate which is a determinant factor facilitating precipitation of these secondary minerals. The solubility of any silica mineral (quartz, chalcedony etc.) can be presented as shown in Equation (2), below.

SiO 2 ( s ) + 2H 2 O ( l ) ⇄ H 4 SiO 4 ( aq ) ⇄ H 3 SiO 4 ( aq ) − + H ( aq ) + (2)

and the solubility constant is given by; KSiO₂ ⇆ aH₄SiO₄.

Silicic acid is a weak acid and dissociates to yield hydrogen ions. If the pH of the solution is increased (becomes more alkaline) then the solubility of silica will also increase as the hydrogen ions are consumed by reaction given in Equation (3).

HCO 3 ( l ) − + H ( aq ) + ⇄ H 2 CO 3 ( aq ) ⇄ H 2 O + CO 2 ( g ) (3)

Deposition of amorphous silica possesses high mechanical strength and is, therefore difficult to remove. This may necessitate an extra cost in maintenance and workover engineering. Highly altered and fractured zones where alteration is dominant are frequently the problematic areas during drilling. These zones are marked by drilling and circulation challenges such as mud loss and pipes getting stuck due to formation slumping (Aldred et al., 1999).

Olkaria geothermal field is located within the Greater Olkaria Volcanic Complex (GOVC) in the South-Central Kenyan Rift Valley, 125 km northwest of Nairobi city, southwest of Lake Naivasha (Figure 1) and along the East African Rift System (EARS) that transects the country from north to south.

To the north the GOVC is bound by Eburru to the north, Longonot to the east, and Suswa to the south (Musonye, 2015). Olkaria geothermal field is a high-temperature geothermal field with temperatures greater than 200˚C. With an estimated resource area of 204 km², Olkaria is divided into seven fields for

management purposes. These include the Olkaria Southeast, Olkaria South west, East Production fields, Olkaria Northeast, North-west field Central fields and the Domes fields (Figure 2).

Each field is at a different stage of development (Okoo, 2013; Ofwona, 2002). Well, OW-38A is found in the Olkaria East production fields while OW-910 and OW-910A are found in the Olkaria Domes.

The East African Rift System, EARS is part of the intra-continental plate divergence rift complex characterised by intense volcanism that took place from late

Tertiary to Recent (Macdonald et al., 1992). The Red Sea and the Gulf of Aden are two arms of a postulated triple junction centred on the Afar region. Oceanic crust began to form under these arms since 10 Ma. The third arm, the EARS, extends southwards through Ethiopia, Kenya and Tanzania into Mozambique. The rifting outcome of the African Plate (Nubian plate) from the Arabian plate forms a classic graben averaging 40 - 80 km wide (Simiyu, 2010). Whilst lithospheric extension is occurring beneath the EARS, the rift has not developed to the stage of ocean crust formation. Along the EARS, most volcanic centres had one or more explosive phase including caldera collapse. The volcanic centres host extensive geothermal systems and are spread with hydrothermal activity.

The rifting in the EARS propagated south and separated into the Eastern and Western branches. The western branch extends over a distance of more than 2100 km from Lake Albert (Mobutu) in the north to Lake Malawi (Nyasa) in the south (Chorowicz, 2005). The eastern arm segment of the rift valley, also called the Kenyan Rift System (KRS), extends from Lake Turkana in Northern Kenya (where the earliest volcanic rocks in the Kenyan Rift have been found) to Lake Natron in northern Tanzania (MacDonald, 2002). The Kenyan rift is centred on the hot spot represented by the Kenya dome uplift, which is elliptical in plan and about 1000 km wide. It has three rift arms, two of which form the main rift, whereas the third is subdued, trending west from the centre of the dome (Okoo, 2013). The geology and geological setting of Olkaria geothermal system is entirely in a major volcanic complex that has been cut by an N-S trending normal rifting faults inside the Kenyan rift (Axelsson et al., 2013). This volcanic complex entails at least 80 smaller volcanic centres containing cementite or peralkaline rhyolites. The structural occurrence of these volcanic centres is either steep-sided domes formed out of lava or pyroclastic rock or as thick lava flows. The rocks within Olkaria geothermal systems are associated with quaternary volcanism and consist of mildly peralkaline silicic volcanic domes, lava flows and the air-fall pumice and peripheral basalts (Omenda, 1998; Lagat et al., 2004). The sub-surface geology of Olkaria consists of a thick volcanic pile of predominantly alkaline, silicic rocks and pyroclastic materials with minor basaltic intercalations (Omenda, 1998).

The structural development in GOVC resulted from the tectonic-thermal rifting and lithospheric thinning during the development of the EARS. The main structures controlling fluid flows include the ring structure, rift fault systems, the Ol’Njorowa gorge, and dyke swarms (Figure 3).

The faults trends are ENE-WSW, N-S, NNE-SSW, NW-SE and WNW-ESE. The NW-SE and WNW-ESE faults are thought to be the oldest fault system and they link the parallel rift basins to the main extensional zone (Wheeler & Karson, 1994). Gorge Farm fault is the most prominent of these faults. It bounds the geothermal fields in the north-eastern part and extends to the Olkaria Domes area. The most recent structures are the N-S (Ololbuttot eruptive fissure) and the NNE-SSW faults (Okoo et al., 2017). Outstandingly, subsurface faults have been encountered in most wells drilled in Olkaria. This was proved by drilling problems such as cave-ins and loss of drilling fluid encountered when these faults were dissected during the drilling process (Lagat, 2009).

Drill sample analysis uncovered, six primary lithological units: pyroclastics, tuff, rhyolite, trachyte, basalt, and intrusive (Figure 4).

The uppermost unit composed of tuff with a matrix of brownish to grey unconsolidated pyroclastics consisting of ash, glass, pumice, obsidian and lithic fragments of weathered lava. This layer is presumed to be part of the upper Olkaria volcanic ranging from a depth of 0 m to 50 m reference to the surface as documented in literature by (Kandie, 2018). Rhyolite was among the dominant strata, rich in silica and composed of mostly light-light brown, greyish-fine grained quartz and feldspar phyric lava. Characterised by moderately altered to greenish and brownish crystalline. Overriding formation in the reservoir rock

was trachyte with quartz and feldspar rich greenish-grey to brownish fine-grained phyric lava. The trachyte displayed moderate to high intensity of alteration to brownish clays. Veined and fractured structures were also observed. Basalt formed the Olkaria geothermal cap rock (Omenda, 1998), which was characterised by dark grey, dark brown fine-grained, highly porphyritic lava with feldspars.

The main secondary minerals identified included oxides, clays, chloride, calcite, quartz, pyrite and epidote. The overall sequence of alteration minerals investigated from the formation in the wells consisted of:

1) Pyroclastics alteration minerals: chalcedony, oxides.

2) Rhyolite alteration minerals: calcite, chalcedony, chlorite, oxides, pyrite.

3) Trachyte alteration minerals: calcite, chalcedony, chlorite, oxides, pyrite.

4) Tuff alteration minerals: calcite, chlorite, mordenite, oxides, pyrite, quartz.

5) Basalt alteration minerals: calcite, chalcedony, chlorite, epidote, illite, oxides, pyrite, quartz.

Intense and irregular pattern of loss of return have been identified in the wells. This scenario in OW-910A (Figure 4) is complemented by abundance of clays and quartz suggested possible interaction of fluids in the veins with the host rock. The presence of clays, quartz and chlorite has been used as evidence of water-rock interaction (Murakami et al., 1996).

The characteristics of secondary minerals including chlorite, epidote, calcite, pyrite and base metal sulphides highlighted an indication of argillic-propylitic alteration (Table 1). This is probably from the decomposition of basalts, rhyolite and trachytes. Epidote may have correspondingly resulted from the replacement of feldspar in silicic rocks given that albite is known to replace feldspars (Reyes, 2000).

The observed silica (amorphous silica and quartz) in the wells during this study occurs at topmost depth of 1018 m to 2410 m in OW-910 and to nearly total depth (TD) in other wells. Manifestation of quartz suggests silicification (addition of secondary silica). This is confirmed by scale deposits markedly observed at wellhead and steam lines in OW-910A and OW-910 respectively (Figure 5).

Dominant occurrences of clays (smectite, illite, chlorite and calcite) was notable. This could be an indication of acidic or near neutral chemical environment in the geothermal reservoir during equilibration of reservoir fluid with the host rock at various reservoir conditions of alkalinity, temperature and pressure (Velde, 1995).

The composition of the equilibrated sample was calculated on basis of the reservoir having boiled and lost steam before sampling, after single-stage adiabatic boiling and computed fluid sample after conductive cooling. The outcome of the data from chemical analyses of samples collected at surface are recorded in (Table 2 and Table 3).

A mineral solution equilibrium is attained when major elements; cation/proton activity ratios achieve a particular value at a specific temperature. The activity ratios may vary with a different composition from basaltic to silicic volcanic samples. Cation/proton activity ratios of the samples in this study have been plotted against temperature (Figure 6).

The reservoir water shows low Ca²⁺/H⁺ ratios to Na⁺/H⁺ and K⁺/H⁺ activity ratios. Na⁺/H⁺ and K⁺/H⁺ activity ratios plot mutually in the equilibrium curve with slight nonconformity. Considering the cation/proton ratios for OW-38A, OW-910 and OW-910A and the fact that K⁺/H⁺ and Na⁺/H⁺ ratios are much higher than the Ca²⁺/H⁺ ratios, aquifer temperatures for the wells under this study are the Na-K geothermometry temperatures of the discharges. This is due to the anomalous concentrations of “indicator” elements in groundwaters and

gases samples collected (Fournier, 1977). The overall decline in the trend of activity ratio of Ca²⁺/H⁺ in the wells with increase in temperature (Figure 6) may be attributed to loss of calcium from the ascending solution in the wells as it boils and precipitates calcite (Tarcan, 2001). Therefore, mixed waters tend to be calcite undersaturated and with low Ca²⁺/H⁺ activity ratios compared with geothermal waters (K⁺ and Na⁺) which decrease the cation/proton activity ratios (Arnórsson, 1985).

These plots were used to estimate heat and silica balance in the reservoirs and balance of silica with other components on assumptions of no heat loss after mixing, quartz solubility controls the silica content of the reservoir fluid and no dissolution and deposition of silica before and after mixing (Figure 7).

Under hydrothermal conditions silica chemistry control phase stability, the trends of enthalpy-silica plots were dispersed between 2000 - 2400 kj/kg and 400 - 800 mg/kg this is due to the solubility of silica in geothermal solutions, silica mixing model and silica geothermometry with adiabatic cooling.

As opposed to enthalpy-SiO₂ plots, Silica-Chloride plots (Figure 8) presents a good linear relationship, presumably mixing with no dissolution of silica from the rock and no precipitation had occurred. The trend changed in areas of low Cl concentration. This might be attributed to dilution devoid of boiling. In any case, if unmixed water was in equilibrium with respect to quartz and re-equilibration of the phase occurred after mixing then, considerate silica would have to precipitate from the water.

High Cl concentration in sample indicates water fed directly from a deep reservoir with minimal mixing or conductive cooling, these are characteristics of groundwater dilution in accordance to (Nicholson, 1993). As observed in Figure 8(a) and Figure 8(b).

The Cl and B are preservative ions and do not take part in chemical reactions in the reservoir (Arnórsson, 1985). Therefore, Cl and B are not incorporated in the alteration minerals in the reservoir. Most importantly, the concentration of Cl is determined by its concentration in the case of mixing geothermal water with cold water. B/Cl ratios presented a good linear relationship between chloride and boron concentrations from these wells, with Cl closely in equilibrium with the reservoir water in OW-38A and OW-910. Linear relationship between chloride and bromide constitutes the main evidence for mixing which is further substantiated by, silica and relations (Arnórsson, 1985). The joint of a regression line through the data points at 0 mg/kg boron and chloride for OW-38A and OW-910 (Figure 8(a)) indicate a common reservoir source for waters. This is reserved to present strong evidence for mixing geothermal water with cold groundwater and gives the impression of boiling preceding mixing.

Characteristics of steam-heated waters were observed in OW-38A and OW-910A owing to the results from plots in Figure 8(c), Figure 8(d) and Figure 9(a) and Figure 9(b). This is noticeable by the low Cl and relatively high sulphate concentrations Figure 9(a) and Figure 9(c). The positive linear correlation between chloride and sulphate may be due to possible boiling and rise in sulphate resulting from increasing oxidation of the hydrogen sulphide as the fluid ascends. Owing to low pH of the steam heated acid-sulphate waters, there are high chances of acid water effectually dissolving the primary minerals of common rocks leaving a residue rich in amorphous silica, native Sulphur, sulphides and aluminous sulphates (Churchman & Lowe, 2012).

The wells presented mixing of CO₂ with deep geothermal water in OW-910A Figure 9(c). The probable origin of CO₂ may have been from the mixing of mantle-derived magmatic or metamorphic CO₂ with ground or surface waters from which it is then steam-heated and diluted with shallower water upon boiling (Lowenstern et al., 2015).

Ideally, volcanic degassing increases the gas concentration in reservoirs, among others including CO₂, H₂S, H₂ and CH₄. The geothermal waters lose part of these gases at the onset of boiling in the reservoir. The reservoir processes considered were boiling and fluid-fluid mixing. The results in the gas concentration of the geothermal fluids in this study were variable with CO₂ being the major constituent at <1% volume, while hydrogen sulphide and hydrogen were relatively abundant. Methane and nitrogen concentrations were variable. Plots of absolute partial gas pressures versus temperatures indicated that the gas pressures increased steadily with temperatures (Figure 10). This is in agreement with kinetic theory with respect to Dalton’s and Raoult’s laws of partial pressures (Othmer et al., 1953).

The state of calcite and amorphous silica saturation was evaluated upon single adiabatic boiling at various temperatures, the computed saturation indices (SI), given by SI = (Log Q/LogK) have been plotted against the temperature of the boiling brine Figure 10. The results from the evaluated wells yielded potential calcite and amorphous silica scaling within the wells. The waters are supersaturated with calcite and amorphous silica (positive saturation Indices). However, amorphous silica exhibits more saturation tendencies than calcite (Figure 11).

The sharp change in slope of saturation index (SI) between 50˚C and 100˚C in OW-910 (Figure 11) may be attributed to the action of degassing of the fluid loss during boiling (Tarcan, 2001). Boiling causes a drastic decrease in CO₂ partial pressures, which leads to calcite oversaturation and precipitation. The oversaturation reaches extreme at a temperature that is a few tens of degrees lower than the temperature at which boiling sets in. At a peak of 100˚C, the water has almost quantitatively been degassed. Continuous boiling, which leads to cooling, will consecutively cause less oversaturation as the solubility of calcite increases with decreasing temperature. In OW-38A, calcite appears to have reached maximum saturation in the fluid and it is evident that decline in solubility is attributed to cooling due to boiling (Simmons & Christenson, 1994).

The solubility of silica can be substantiated at a higher temperature, strong conductive cooling and mixing in up-flow zones may produce undersaturated fluid with a relatively high silica and Na⁺/K⁺ ratio. Therefore, amorphous silica precipitation is related to water temperature, the degree of water oversaturation and water salinity (Arnórsson et al., 2007b). Consequently, the cooling, which results from boiling, causes the water in the wells to become oversaturated with respect to amorphous silica because its solubility decreases with decreasing temperature. The trend in Figure 11 indicates amorphous silica is convenient for geothermometric purposes (Arnórsson & Gunnlaugsson, 1985).