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Geothermal systems in the Sunda volcanic island arc
Investigations on the islands of Java and Bali,
Indonesia
Dissertation
zur Erlangung des Doktorgrads der Naturwissenschaften (Dr. rer. nat.)
am Fachbereich Geowissenschaften
Universität Bremen
vorgelegt von
Budi Joko Purnomo
Bremen
March, 2015
ii
iii
Gutachter:
Prof. Dr. Thomas Pichler
Prof. Dr. Peter La Femina
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v
E r k l ä r u n g
Hiermit versichere ich, dass ich
i. die Arbeit ohne unerlaubte fremde Hilfe angefertigt habe,
ii. keine anderen als die von mir angegebenen Quellen und Hilfsmittel
benutzt haben und
iii. die den benutzten Werken wörtlich oder inhaltlich entnommen Stellen
als solche kenntlich gemacht habe.
___________________ ,den ________________
______________________________
(Unterschrift)
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TABLE OF CONTENTS
Abstract .......................................................................................................................... ix
Zusammenfassung ....................................................................................................... xi
I. Introduction ................................................................................................................ 1
I.1. Thesis outline ................................................................................................................. 2
I.2. Research objectives ...................................................................................................... 4
I.3. Theoretical backgrounds .............................................................................................. 5
I.3.1. An overview of geothermal system ...................................................................... 5
I.3.2. Geothermal manifestations ................................................................................... 8
I.3.3. Geothermometry ..................................................................................................... 9
I.3.4. Boron isotope ........................................................................................................ 14
II. Geological setting, sampling and analyses ......................................................... 17
II.1. Geological setting ....................................................................................................... 17
II.2. Sampling and analyses .............................................................................................. 20
III. Geothermal systems on the island of Java, Indonesia ..................................... 23
Abstract ................................................................................................................................ 24
III.1. Introduction ................................................................................................................. 25
III.2. Sampling locations .................................................................................................... 25
III.3. Results ........................................................................................................................ 26
III.4. Discussion .................................................................................................................. 30
III.4.1. General considerations about geothermal systems on Java....................... 30
III.4.2. Geothermometry................................................................................................. 37
III.4.3. The heat sources of the fault-hosted geothermal systems .......................... 41
III.4.4. Oxygen and hydrogen isotope considerations .............................................. 44
III.5. Conclusions ................................................................................................................ 46
IV. Boron isotope variations in geothermal systems on Java, Indonesia ............49
Abstract ................................................................................................................................ 50
IV.1. Introduction ................................................................................................................ 51
IV.2. Sampling locations .................................................................................................... 52
viii
IV.3. Results ........................................................................................................................ 52
IV.4. Discussion .................................................................................................................. 56
IV.4.1. Boron in thermal waters and seawater input ................................................. 56
IV.4.2. The δ11B of acid sulfate and acid chloride crater lakes ................................ 60
IV.4.3. Processes affecting the δ11B value of thermal waters ................................. 62
IV.4.4. The δ11B of fault-hosted and volcano-hosted geothermal systems ........... 66
IV.5. Conclusions ................................................................................................................ 66
V. Geothermal systems on the island of Bali, Indonesia ........................................ 69
Abstract ................................................................................................................................ 70
V.1. Introduction.................................................................................................................. 71
V.2. Geological setting ....................................................................................................... 72
V.3. Results ......................................................................................................................... 74
V.4. Discussion ................................................................................................................... 75
V.4.1. Geochemistry of thermal waters ....................................................................... 75
V.4.2. Phase separation and seawater input ............................................................. 82
V.4.3. Geothermometry ................................................................................................. 84
V.4.4. Oxygen and hydrogen isotope considerations ............................................... 86
V.5. Conclusions ................................................................................................................. 88
VI. Conclusions and outlook ...................................................................................... 91
VI.1. Conclusions ................................................................................................................ 91
VI.2. Outlook........................................................................................................................ 92
Acknowledgements ..................................................................................................... 95
References .................................................................................................................... 97
ix
Abstract Geothermal systems located in a volcanic island arc are generally considered to be
exclusively influenced by Quaternary volcanic activities. Giving the complex geological
setting on the island of Java, the geothermal systems probably is not that simple. The
subduction of the Indo-Australian plate beneath the Eurasian plate formed three
volcanisms, Paleogene, Neogene and Quaternary, and faults. The subduction also
thinned the crust of the southern part of the island due to uplift and erosion. The
presence of geothermal systems hosted by a volcano and a fault zone, thus named as
volcano-hosted and fault-hosted, respectively, give an opportunity to study their
different physicochemical characteristics. The influence of Quaternary magma in the
fault-hosted geothermal systems distributed in two different volcanic belts, Quaternary
and Tertiary, were examined. Additionally, the contrast signatures of boron isotope
between two different crater lakes, acid sulfate and acid chloride, which likely has
been overlooked, were determined. Finally, the possibility of a carbonate sedimentary
basement as the host-rock of geothermal systems on the volcanic island of Bali was
investigated.
The physicochemical properties of geothermal waters were used in this study,
including major anion and cation, trace elements and stable isotope (2H and 18O). The
fractionation characteristics of boron isotope were used for further investigation on the
contrasting fault-hosted and volcano-hosted geothermal systems. The distinct boron
isotope composition of seawater was applied to confirm seawater input in some
geothermal systems.
The volcano-hosted and fault-hosted geothermal systems were chemically different:
the former had higher HCO3- concentrations and Mg/Na ratios compared to the latter.
This condition was caused by CO2 magmatic gas supply in the volcano-hosted, which
was insignificant or absent in the fault-hosted geothermal systems. The CO2 gas
supply produces slightly acid HCO3- thermal waters, hence together with its slower
ascent due to the longer flow path would elevate the Mg2+ content. Geothermal
systems hosted by faults located in the Quaternary magmatic belt were clearly
supplied by magmatic fluid, thus could not be classified as a fault-hosted geothermal
system. Although the Quaternary magmatic fluid input in the fault-hosted geothermal
systems located in the Tertiary volcanic belt are absent, the reservoir temperature and
lithium (Li) enrichment indicated a Quaternary magmatic heat source for the fault-
x
hosted geothermal systems of Cilayu and Cisolok. Meanwhile, the other fault-hosted
geothermal systems are likely heated by a deep-seated magma. Magmatic fluid input
enriched the 2H and 18O isotope of some volcano-hosted geothermal systems,
something that was not identified in any of the fault-hosted geothermal systems.
Boron isotope further distinguished the volcano-hosted and fault-hosted geothermal
systems. The magmatic fluids input, favorable for minerals precipitation, and longer
thermal water ascent in the former geothermal system promoted δ11B enrichment. In
contrast, the fast thermal water ascent and absence of magmatic fluids input
maintained a light δ11B signature in the fault-hosted geothermal systems. Two
different types of thermal crater lakes, acid sulfate and acid chloride, had a contrast in
B isotope compositions. The acid chloride crater lake had a light δ11B value
representing magmatic origin, while the acid sulfate crater lakes had heavier δ11B
values, produced by a sequence of processes: vapor phase separation in the
subsurface, followed by evaporation and B adsorption into clay minerals on the
surface. B isotope confirmed seawater input in two fault-hosted geothermal systems:
Parangtritis and Krakal.
The geothermal systems on Bali were studied using the physicochemical properties of
surface thermal waters. The (Ca2+ + Mg2+)/HCO3- of approximately 0.4 and the visual
absence of 18O isotope enrichment ruled out a carbonate host rock type, instead the
K/Mg ratios indicated water-rock interaction with a calc-alkaline magmatic rocks. The
well correlation of HCO3- content with Ca2+, Mg2+, Sr2+ and K+ revealed water-rock
interaction influenced by carbonic acid. The B/Cl ratios revealed phase separation for
the Bedugul and Banjar geothermal systems. The heavy δ11B of +22.5 ‰ and a Cl/B
ratio of 820 confirmed seawater input in the Banyuwedang geothermal system.
xi
Zusammenfassung Es wird allgemein angenommen, dass Geothermale Systeme in vulkanischen
Inselbögen ausschließlich durch vulkanische Aktivität angetrieben werden. Die
komplexe geologische Struktur der Insel Java legt nahe, das dieses Erklärungsmodell
dort wahrscheinlich zu einfach ist. Durch Subduktion der indo-australischen unter die
eurasische Platte verursachter Vulkanismus, sowie dadurch entstandene
Grabensysteme, lassen sich chronostratigraphisch dem Paläogen, Neogen und
Quartär zuordnen. Die Mächtigkeit der Kruste ist durch Heraushebung und Erosion
während des Subduktionsprozesses verringert. Die geographische Nähe
geothermaler Systeme die sich an Vulkanen („volcano-hosted“) oder in Grabenbrüche
(„fault-hosted“) befinden, erlaubt es, die unterschiedlichen physiko-chemischen
Eigenschaften beider Systeme zu untersuchen. In den vorliegenden Arbeiten wurden
die Auswirkungen quartären Magmas auf “fault-hosted” geothermale Systeme
untersucht. In zwei Kraterseen wurde die Isotopenverteilung des Elements Bor (δ11B)
verglichen und die Konzentrationen von gelöstem Sulfat und Chlorid, die in vorherigen
Arbeiten häufig nicht betrachtet werden, bestimmt. Zusätzlich wurde untersucht, ob
karbonatisches Sediment als Untergrund in geothermalen Systemen der vulkanisch
entstandenen Insel Bali möglich ist.
Die physiko-chemischen Eigenschaften geothermalen Wassers, einschließlich
Anionen- und Kationenverteilungen, Konzentrationen von Spurenelementen sowie
Isotopengehalte (δ2H, δ18O) wurden untersucht. Die Isotopenverteilung des Elements
Bor (δ11B) wurde zur weiteren Charakterisierung der Unterschiede von „volcano-
hosted“ und „fault-hosted“ geothermalen Systemen genutzt. Die bekannte
Konzentration des δ11B in Meerwasser wurde zusätzlich genutzt, um das Einströmen
von Meerwasser in einige geothermale Systeme zu bestätigen.
„Volcano-hosted“ und „fault-hosted“ geothermale Systeme lassen sich aufgrund ihrer
chemischen Eigenschaften unterscheiden. Erstere haben höhere HCO3-
Konzentrationen und Mg2+/Na+-Verhältnisse verglichen mit Letzteren. Dieser Befund
lässt sich auf den Einfluss des CO2-Gehalts in magmatischen Gasen zurückführen,
der in “fault-hosted” Systemen vernachlässigbar gering bzw. nicht vorhanden ist.
Durch CO2-Eintrag wird in geothermalem „volcano-hosted“ Systemen
Hydrogenkarbonat gebildet und durch das dissoziierende Proton eine Senkung des
pH-Werts verursacht. Aufgrund der, durch die geologische Architektur bedingten,
xii
verlängerten Passage durch das Gestein erhöht sich der Mg2+-Gehalt. Die durch
Grabenbildung in quartären Vulkangürteln entstandenen geothermalen Systeme
werden eindeutig durch Wärmekonvektion magmatischer Flüssigkeit gespeist, was
dazu führt, dass diese nicht als „fault-hosted“ geothermales System im eigentlichen
Sinn beschrieben werden können. Obwohl ein Einströmen Quartär gebildeten
Magmas in die “fault-hosted” geothermalen Systeme, Cilayu und Cisolok, die in einem
im tertiär gebildeten Vulkangürtel lokalisiert sind, nicht gezeigt werden kann, kann
durch die Reservoir-Temperatur sowie die Anreicherung von Lithium (Li) Quartär
entstandenes Magma als Wärmequelle angenommen werden. Wahrscheinlich ist der
Ursprung anderer “fault-hosted” Systeme dieser Region die Wärme tiefliegender
Magmakörper. Durch magmatischen Einfluss werden 2H- und 18O-Isotope im Wasser
von “volcano-hosted” Systemen angereichert. Dieses Phänomen konnte in keinem
der “fault-hosted” Systeme beobachtet werden.
Eine weitere Möglichkeit der Unterscheidung geothermaler Systeme zwischen
“volcano-hosted” und “fault-hosted“ besteht in der Messung des δ11B. Durch
magmatische Flüssigkeit wird die Ausfällung von Mineralien unterstützt, und die oben
beschriebene längere Passage führt zu einer höheren Anreicherung von δ11B. Im
Gegensatz dazu führt eine schnelle Passage ohne Einströmen magmatischer
Flüssigkeit nur zu einer abgeschwächten Konzentration von δ11B in „fault-hosted“
geothermalen Systemen. Durch geothermale Systeme entstandene Kraterseen
können aufgrund ihrer δ11B-Konzetration der Genese charakterisiert werden.
Kraterseen mit hoher Cl--Konzentration haben einen niedrigen δ11B-Wert und lassen
auf magmatischen Ursprung schließen, während Kraterseen mit einer hohen Sulfat-
Konzentration ebenfalls einen hohen δ11B-Wert aufweisen aber im wesentlichen mit
Wasser meteorischen Ursprungs gefüllt sind. Dem hohen δ11B-Wert liegt der Ablauf
folgender Prozesse zugrunde: unterhalb der Oberfläche kommt es zu einer Trennung
der Dampfphase, durch anschließende Verdunstung adsorbiert Bor an Tonmineralien
der Oberfläche. Die hohen δ11B Werte bestätigen den Meerwassereintrag in die „fault-
hosted“ geothermalen Systeme von Parangtritis und Krakal.
Geothermale Systeme der Insel Bali wurden anhand der physiko-chemischen
Eigenschaften des geothermalen Oberflächenwassers untersucht. Ein ([Ca2+]+
[Mg2+])/HCO3- -Verhältnis von ~0,4 und eine nicht vorhandene Anreicherung von δ18O
schließen den Einfluss von Karbonat-Sediment aus. Die in den Untersuchungen
bestimmten K+/Mg2+-Verhältnisse deuten hingegen auf eine Einbeziehung von
xiii
basischer Magma in den Reaktionskreislauf. Die beobachteten hohen Korrelationen
der HCO3- Gehalte mit Ca2+, Mg2+, Sr2+ und K+ lassen auf eine Kohlensäure vermittelte
Reaktion an der Phasengrenze Wasser-Stein schließen. Eine Phasentrennung in den
geothermalen Systemen von Bedugul und Banjar konnten durch die B/Cl-Ratios
gezeigt werden. Ein δ11B von +22.5‰ und ein Cl/B-Verhältnis von 820 bestätigten
den Meerwassereintrag in das geothermale System von Banyuwedang.
xiv
INTRODUCTION
1
I. Introduction Indonesia is known for having the biggest geothermal resources in the world,
mostly spread along the Sunda volcanic island arc, which extends from west to east,
from the island of Sumatera to the island of Damar. The island of Java located in the
middle of the island arc is the most geothermal exploited area due to the huge
demand of electricity. Based solely on the geological association, hosted by either a
volcano or fault, geothermal systems on Java can be divided into volcano-hosted and
fault-hosted, respectively. Concerning their potential as an electricity power source, to
date the geothermal development program has been only focused on the volcano-
hosted geothermal systems due to the assumption of low potential energy in the fault-
hosted geothermal systems.
Apart from the energy issue, the presence of fault-hosted geothermal systems
on Java provides an essential opportunity to determine the influence of the
Quaternary volcanoes to the systems. In a volcanic islands arc, the geothermal
systems are considered to be simply dominated by the Quaternary volcanic activities.
Therefore, less attention is given to the fault-hosted geothermal system located in this
geological setting. Several fault-hosted geothermal systems distributed in a
Quaternary volcanic arc, for instances in the Liquine-Ofqui fault zone, Chile and in the
Southern Apennines, Italy (Alam et al., 2010; Italiano et al., 2010), indicated
contrasting geochemical processes compared to volcano-hosted geothermal systems.
In the fault-hosted system deep circulated groundwater is simply conductive heated,
while in the volcano-hosted system involves condensation of volcanic steam (Alam et
al., 2010). On Java the geothermal systems are probably more complex due to the
geological setting. The island is mainly developed by three main volcanisms:
Paleogene volcanism formed the Tertiary volcanic belt, distributed along the southern
part, followed by Neogene volcanism shifted northward to the middle of the island and
finally Quaternary volcanism emerged along the Neogene volcanic (Soeria-Atmadja et
al., 1994). The relatively impermeable Tertiary volcanic rocks might inhibit the extent
of Quaternary magmatic fluids flow. The subduction in the south of Java developed
several major and minor faults that are distributed in the two volcanic belts, Tertiary
and Quaternary. The northward plates subduction has uplifted the southern part of the
island, followed by erosion, hence produced a thinner crust (Clements et al., 2009;
Hall et al., 2007). The thin crust enables deep penetrated groundwater through fault to
extract the heat. Considering these geological settings, different characteristics
INTRODUCTION
2
between fault-hosted located in the Quaternary volcanic belt and the Tertiary volcanic
belt could be expected.
Two characteristics of boron (B) isotope fractionation might be beneficial for
further investigation on the contrast between volcano-hosted and fault-hosted
geothermal systems. Those characteristics are 1) the relatively conservative at
temperatures >65 °C (Aggarwal and Palmer, 1995; Palmer et al., 1987) and 2) the
significant fractionation due to B adsorption/coprecipitation onto clay, iron oxide and
calcite and evaporate minerals (Agyei and McMullen, 1968; Hemming and Hanson,
1992; Lemarchand et al., 2007; Oi et al., 1989; Palmer et al., 1987; Schwarcz et al.,
1969; Spivack and Edmond, 1987; Swihart et al., 1986; Vengosh et al., 1991a;
Vengosh et al., 1991b; Vengosh et al., 1992). The fast ascent of fault-hosted thermal
water through a fault and the absence of magmatic fluids input should result in a light
δ11B signature. Conversely the slow ascent and the presence of magmatic fluids input
might be favorable for δ11B enrichment. Additionally the presence of two contrasting
thermal crater lakes, acid-sulfate and acid-chloride, on Java provide an opportunity to
determine their distinct boron isotope signatures, which likely has been overlooked so
far.
Finally, the study on the characteristic of volcano-hosted and fault-hosted
geothermal systems was extended to the volcanic island of Bali. Apart from this
objective, the island is an ideal location to investigate the role of a carbonate rock
basement to the geothermal systems. In such a geological setting, the carbonate rock
was identified hosting the geothermal systems instead of volcanic rocks, e.g., in
Vicano-Cimino and Sabatini-Tolfa (Cinti et al., 2011; Cinti et al., 2014).
I.1. Thesis outline This thesis consists of six chapters with three main bodies. Chapter 1,
Introduction, provides the context of the research and overview of some theoretical
backgrounds. Chapter 2 explains the geological setting of the study areas, sampling
and analytical methods. Chapter 6 presents the overall conclusion answering the
research objectives and outlook related to the results of this study. Below the three
main chapters and my contributions on every chapter are briefly described.
Chapter 3
This chapter provides a classification of the geothermal systems on Java. The
systems were classified into volcano-hosted and fault-hosted based solely on their
INTRODUCTION
3
geological association, either in a volcano or fault zone. Subsequently the
physicochemistry of thermal waters clearly divided the two geothermal systems. The
possible role of Quaternary magma as a heat source for fault-hosted geothermal
system was investigated based on the reservoir temperatures and trace element
enrichments.
This research is published in Journal of Volcanology and Geothermal Research
(Purnomo, B.J. & Pichler, T., 2014).
My contributions on this article are:
- sampling and measurement of some physicochemical parameters in the field,
- preparing samples for laboratory analyses,
- quality control on the laboratory results,
- data interpretation, and
- writing the first draft of the published paper.
Chapter 4
In this chapter boron isotope was used to further investigate the distinct
geochemical characteristics between volcano-hosted and fault-hosted geothermal
systems. The contrast in B isotope signatures between two distinct thermal crater
lakes, acid-sulfate and acid-chloride, was determined. Boron isotope was also applied
to confirm seawater input in some geothermal systems on Java.
A part of the result was presented as a poster (Purnomo, B.J., Pichler, T. and
You, C.-F, 2014) in American Geophysical Union (AGU) Fall Meeting 2014. The final
version of this chapter has been submitted to Geochimica et Cosmochimica Acta
(Purnomo, B.J., Pichler, T. and You, C.-F, submitted in November 2014, current
status: under review).
My contributions on this article are:
- sampling and measurement of some physicochemical parameters in the field,
- preparing samples for laboratory analyses,
- data interpretation, and
- writing the first draft of the submitted manuscript.
Chapter 5
This chapter presents an overview of geothermal systems on Bali that were
classified into a volcano-hosted and fault-hosted. However, the main focus is to
investigate the host-rock, either carbonate or volcanic, considering the Quaternary
INTRODUCTION
4
volcanoes have a carbonate sedimentary basement. The processes that govern the
chemical composition of the surface thermal waters, seawater input and
geothermometry of the reservoir were also determined.
The result of this study has been submitted to Journal of Volcanology and
Geothermal Research (Purnomo, B.J. and Pichler, T., submitted in February 2015,
current status: under review).
My contributions on this article are:
- sampling and measurement of some physicochemical parameters in the field,
- preparing samples for laboratory analyses,
- quality control on the laboratory results,
- data interpretation, and
- writing the first draft of the submitted manuscript.
I.2. Research objectives The main purpose of this study is to characterize the system of geothermal
fields located in the Sunda volcanic island arc. The geothermal system probably is
complicated by the complex geological setting formed by the dynamic plates
boundary. The presence of multiple volcanisms, fault zones and thin crust potentially
influence the geothermal system, instead of exclusively dominated by Quaternary
volcanic activities. Several specific objectives of the study are:
to define the geothermal systems into a volcano-hosted or fault-hosted systems,
to calculate the reservoir temperatures of the geothermal systems,
to characterize the contrast between volcano-hosted and fault-hosted
geothermal systems, including the physicochemical processes and fluids origin,
to investigate the role of Quaternary volcanoes to the fault-hosted geothermal
systems,
to determine the possible heat sources of the fault-hosted geothermal systems,
to apply boron isotope for further investigation on the contrast between fault-
hosted and volcano-hosted geothermal systems,
to investigate the boron isotope signatures of two distinct crater lakes, acid-
sulfate and acid-chloride,
to confirm seawater input in some geothermal systems, considering the location
close to sea, and
INTRODUCTION
5
to determine the role of a carbonate sedimentary basement on the volcanic
island of Bali.
I.3. Theoretical backgrounds I.3.1. An overview of geothermal system
Geothermal system is a heat transfer process from the Earth’s crust to the
surface. Generally the heat transfer involve meteoric water, hence it is called as
hydrothermal system (Hochstein and Browne, 2000). A geothermal system consists of
three main components: 1) heat source, 2) reservoir and 3) water (e.g., Goff and
Janik, 2000; Kuhn, 2004). The heat sources for geothermal systems include magmas
within the crust, intracrustal nonmagmatic and conductive heat flow (Hochstein and
Browne, 2000). The first can present as a convective magma fluid input or conductive
heat transfer from cooling magma, the second is present due to the anomaly thermal
gradient in areas with relatively thin crust and the third is formed in the sedimentary
basin where thick sediments generate heat and overpressure (Goff and Janik, 2000;
Kuhn, 2004; Nicholson, 1993). Faulds et al. (2010) introduced a term of ‘amagmatic’
heat source for a deep-seated magma, to distinguish it with Quaternary magma.
Apart from those heat sources, back to the experiment by Byerlee (1978), a frictional
heat source was proposed. However, its presence in natural geothermal systems has
never been proved, hence it is referred to ‘the stress-heat flow paradox’ (Lachenbruch
and Sass, 1980).
Several classifications of geothermal systems have been proposed. Simply
based on the reservoir temperature geothermal system was classified into three
(Hochstein and Browne, 2000):
1) High temperature (T >225 °C),
2) Intermediate temperature (125 to 225 °C), and
3) Low temperature (T <125 °C)
However the range of temperatures in the classification is not rigid, for an instance
Benderitter and Cormy (1990) suggested a low temperature geothermal system
accounted for temperatures below 100 °C.
Considering the geology, hydrology and engineering, Goff and Janik (2000)
classified geothermal system into:
1) Young igneous systems
INTRODUCTION
6
Geothermal system hosted by Quaternary volcanic or intrusion. This system
generally spreads along plate boundaries and spreading centers.
2) Tectonic systems
Geothermal system associated with anomaly heat flows due to the relatively thin
crust. The heat is extracted by a deep penetrated groundwater hosted by faults.
3) Geopressured systems
This geothermal system distributed in sedimentary basins. Subsidences and thick
sediments form the heat and overpressure.
4) Hot dry rock systems
The heat stored in an impermeable rock is extracted by injecting water to the rock,
subsequently the resulting thermal water is pumped out to the surface.
5) Magma tap systems
The geothermal system taps the heat of a shallow magma by circulating water.
Nicholson (1993) proposed a classification of geothermal systems based on:
heat transfer, temperature, fluid composition and topographic (Fig. 1.1). Heat transfer
divides geothermal system into two, dynamic and static systems. The former involves
circulation of fluids from depth, where meteoric water circulated into the depth, extract
heat, then ascend to the surface. Meanwhile, in the latter system the thermal fluid is
trapped in the deep strata. The dynamic geothermal systems are distributed in plate
boundaries and spreading centers, while the static system are located in tectonic
stabile areas, for instances in Russia, Eastern Europe and Australia. Considering the
geological setting of the study areas is a subduction zone, only the dynamic system is
further elaborated. The dynamic geothermal system is divided into a high temperature
system, generally associated with Quaternary magmas, and a low temperature
system, associated with conductive heat transfer of anomaly heat flow. The high
temperature system is classified into a liquid-dominated and a vapor-dominated based
on the dominant phase in the reservoir. The contrast topographic of the liquid-
dominated geothermal systems divides the system into a low-relief and a high relief
system. The former is hosted by a silicic volcano, while the latter by an andesitic
volcano, which is a typical of volcanic island arc.
INTRODUCTION
7
Fig. 1.1. Geothermal system classification by Nicholson (1993).
Exclusively based on the geological setting, Saemundsson (2009) classified
geothermal systems into:
1) Volcanic-geothermal system
Geothermal system associated with volcanoes, heated by intrusion or magma.
2) Convective system
Geothermal system generated due to the deep circulation of meteoric water
through fractures, extracted heat from an anomaly high temperature gradient.
3) Sedimentary system
Geothermal system hosted by thick sedimentary layers, heated by regional
thermal gradient.
Working in a complex of major fault and Quaternary volcanic arc geological
setting in the Liquiñe-Ofqui Fault Zone (LOFZ), Alam et al. (2010) divided the
Geothermal system
dynamic (convective)
static (conductive)
high temperature
liquid- dominated
low-relief high-relief
low temperature
geopressured low temperature
vapor- dominated
INTRODUCTION
8
geothermal systems into a volcanic type, associated with stratovolcanoes, and a
structural (non-volcanic) type, hosted by the faults. This classification is adopted in
this study, however, the terms of volcano-hosted and fault-hosted geothermal systems
are preferred. Volcano rather than volcanic is used to avoid confusion with volcanic as
a rock type, while fault is applied to clarify its geological structures.
Although a volcanic arc is generally dominated by volcano-hosted geothermal
systems, fault-hosted geothermal systems might be present, for instances in the
Liquine-Ofqui fault zone, Chile and in the Southern Apennines, Italy (Alam et al.,
2010; Italiano et al., 2010). In the fault-hosted geothermal system a deep circulated
meteoric water through faults is conductive heated, while the volcano-hosted
geothermal system involves condensation of volcanic steam (Alam et al., 2010). The
heat sources of fault-hosted geothermal systems could be a deep seated magma.
This was named as ‘amagmatic’ heat source to distinguish it with shallow magmas of
Quaternary volcanoes, as has been applied in the Great Basin, USA and Western
Turkey (Faulds et al., 2010).
I.3.2. Geothermal manifestations
The existence of a geothermal system in the subsurface can be recognized by
the surface features (manifestation), which include solfatara, fumarole, acid crater
lake, hot spring, mud pool, steaming ground and altered ground (Fournier, 1989;
Giggenbach, 1988; Giggenbach and Stewart, 1982; Goff and Janik, 2000;
Hedenquist, 1990; Henley and Ellis, 1983; Hochstein and Browne, 2000; Nicholson,
1993). These manifestations are formed by a sequence of processes that briefly can
be explained as follow (Giggenbach, 1988; Nicholson, 1993). Volatile magmas
contain gases, H2O, CO2, H2S, HCl and HF, rise and react with deep circulated
groundwater producing acid water. The acid water-rock interaction subsequently
consumes H+ to form a neutral-chloride water, which then ascend to the surface.
During the ascent separation of fluid phases might occur due to a rapid drop in
temperature and pressure. The remaining chloride water discharges as a neutral
chloride hot spring on the surface. In a low relief topography the water emerges above
the upflow zone, close to the acid sulfate and bicarbonate hot springs, while in the
high relief topography (stratovolcano) it discharges far away from the upflow zone to
the flank of stratovolcano (Fig. 1.2). Meanwhile, the vapor phase resulted by the
phase separation usually is rich in CO2 and H2S gases, hence the condensation and
reaction with O2-rich groundwater produce a slightly acid bicarbonate (HCO3-) water
INTRODUCTION
9
and an acid sulfate (SO4-) water, respectively. In a vapor-dominated geothermal
system, thermal water ascends exclusively as a steam phase that rich in CO2 and H2S
gases, therefore, the resulting manifestations are mud pools, acid sulfate and
bicarbonate hot springs (Fig. 1.3).
A geothermal system generally produces three types of hot springs: 1) neutral
chloride, 2) acid sulfate and 3) bicarbonate waters, however, mixtures between the
individual groups are common (Hedenquist, 1990; Hochstein and Browne, 2000;
Nicholson, 1993; White, 1957). The composition of these thermal springs is controlled
by two main factors: 1) reservoir condition and 2) secondary processes during ascent.
Host-rock type, temperature, residence time and magmatic fluid input govern the
geothermal brines. The physicochemistry of thermal waters might change during the
ascent due to phase separation, minerals precipitation, adsorption/desorption, water-
rock interaction, groundwater dilution and seawater input (in a coastal area). The
chemical composition of a thermal spring records the physicochemical processes in
the subsurface. Chemically inert constituent (tracers), Cl, B, Li, Rb and Cs, can be
used to track the source, while chemically reactive species (geoindicators), e.g., Na,
K, Mg, Ca and SiO2, record the physicochemical processes during thermal water
ascent (Ellis and Mahon, 1977; Giggenbach, 1991; Nicholson, 1993).
I.3.3. Geothermometry Solute geothermometers are commonly applied to estimate the reservoir
temperature of a geothermal system, by using either the absolute concentration,
solute ratio or solute relation. The geothermometers were formulated based on the
temperature dependency of a mineral-fluid reaction in the reservoir. This is preserved
on the surface due to the slow re-equilibrium of the minerals in a low temperature
system (Fournier, 1977). The successful application of solute geothermometers for
hot springs relies on five basic assumptions: 1) exclusively temperature dependent
mineral-fluis reaction; 2) abundance the mineral and/or solute; 3) equilibrium reaction
in the reservoir; 4) no re-equilibrium; and 5) no mixing or dilution (Ellis, 1979; Fournier,
1977; Fournier and Truesdell, 1973; Nicholson, 1993; Truesdell, 1975). The last
assumption can be overcome if the extent of dilution can be calculated. Below, solute
geothermometers that were used in this study are listed.
Fig.
1.2
. Con
cept
ual m
odel
of l
iqui
d-do
min
ated
geo
ther
mal
sys
tem
in h
igh
relie
f top
ogra
phic
(stra
tovo
lcan
o) (a
dopt
ed fr
om
H
enle
y an
d E
llis, 1
983)
.
INTRODUCTION
11
Fig. 1.3. Conceptual model of vapor-dominated geothermal system (adopted from Nicholson, 2000).
a) Silica geothermometer
Silica geothermometer is deducted from the solubility of quartz in certain
temperatures and pressures. The thermometer is based on the absolute silica
concentration and thus sensitive to secondary processes such as mixing, boiling and
precipitation (Fournier and Rowe, 1966). Several silica geothermometers were
established for different conditions, such as:
- No steam loss T (°C) = {1309 / (5.19 – log SiO2)} – 273 (Fournier, 1977)
- Max. steam loss T (°C) = {1522 / (5.75 – log SiO2)} – 273 (Fournier, 1977)
In low temperatures the silica content is attributed to the solubility of chalcedony,
cristobalite and amorphous silica (Arnorsson, 1970; Arnorsson, 1975; Fournier and
INTRODUCTION
12
Rowe, 1962; Fournier and Rowe, 1966). Therefore, the silica geothermometer, for an
example chalcedony, is stated as:
- Chalcedony T (°C) = {1032 / (4.69 – log SiO2)} – 273 (Fournier, 1977)
The application of silica geothermometers for hot springs usually calculates
underestimation temperatures, close to the discharge temperature, due to dilution by
shallow water. This can be overcome by calculating the silica ‘parent’ using the silica
mixing model of Truesdell and Fournier (1977) (Fig. 1.4).
Fig. 1.4. Silica vs. enthalpy diagram for silica parent calculation based on Truesdell and Fournier (1977).
b) Na/K geothermometer
The Na/K geothermometer was calculated using the Na2+ and K+ ratios. The
thermometer was formulated based on the temperature-dependent reaction of ion
exchange between albite and K feldspar (Fournier, 1979; Giggenbach, 1988).
NaAlSi3O8 + K+ (aq) = KAlSi3O8 + Na+ (aq) (albite) (K feldspar)
INTRODUCTION
13
The Na/K gethermometer formula for instances are:
T (°C) = {1217 / (log (Na/K) + 1.483} – 273 (Fournier, 1979)
T (°C) = {1390/ (log (Na/K) + 1.750} – 273 (Giggenbach, 1988)
This geothermometer is suitable for reservoir temperatures ranging from 180 to 350
°C, but not for temperatures below 120 °C (Ellis, 1979; Nicholson, 1993). However,
the geothermometer has been indicated resulting in overestimation temperatures
when applied to thermal waters with high Ca2+ concentrations. This is considered due
to the competition of Ca2+, Na+ and K+ during ion exchange reaction (Nicholson,
1993).
c) Na/K/Ca geothermometer
The empiric Na/K/Ca geothermometer was developed by Fournier and Truesdell
(1973) for Ca-rich thermal waters, i.e., (√Ca / Na) >1. The thermometer was defined
as:
T (°C) = 1647 / {log (Na/K) + β[log(√Ca / Na) + 2.06] + 2.47} – 273
The thermometer is considered working well for reservoir temperatures above 180 °C.
However, the application to hot springs could be troublesome. During thermal water
ascent the Ca2+ content might be depleted by CO2 release, Mg2+ exchange and calcite
precipitation.
d) Na/K/Mg geothermometer
Giggenbach (1988) developed a ternary diagram of K/100-√Mg-Na/1000 as a
geothermometer and a tool to assess suitable hot springs for geothermometer
application (Fig. 1.5). In the diagram thermal waters are divided into three groups, fully
equilibrium, partially equilibrium and immature waters. Immature water potentially
produce inaccurate reservoir temperatures calculation.
e) Na/Li geothermometer
The Na/Li ratios of geothermal water were identified having an inverse correlation with
temperatures (Ellis and Wilson, 1960; Koga, 1970). The thermometer was formulated
based on the theoretical exchange reaction of Fouillac and Michard (1981):
Clay-Li + H+ = Clay-H + Li+
INTRODUCTION
14
Fig. 1.5. Evaluation of Na-K-Mg geothermometer (Giggenbach, 1991).
with an equation:
T (°C) = 1000 / {log (Na/Li) + 0.389] – 273 (Fouillac and Michard, 1981)
Later, a new formula was established by Kharaka et al. (1982) as follow:
T (°C) = 1590 / {log (Na/Li) + 0.779] – 273
This thermometer was considered resulting the most reliable reservoir temperatures
for geothermal systems with a carbonate reservoir type (Minissale and Duchi, 1988).
I.3.4. Boron isotope Boron is a trace element, thus can be used to track the thermal water origin.
Dissolve boron is mainly present as B(OH)3 (boric acid, trigonal species) and B(OH)4
(borate anion, tetrahedral species) (Dickson, 1990; Hershey et al., 1986). At low pH
(<7) only B(OH)3 is present, and conversely at pH >10 boron is found as a B(OH)4
INTRODUCTION
15
species. Boron has two stable isotopes, 10B and 11B, with an abundance of 19.8 ‰
and 80.2 ‰, respectively (e.g., Xiao et al., 2013; Bart, 1993). Boron isotope
composition is reported as δ11B per mil (‰) relative to the standard of NIST-SRM 951
(Catanzaro et al., 1970), with a formula:
δ11B (‰) = [(11B/10B)sample / (11B/10B)NIST-SRM-951 – 1] x 1000
Geothermal waters have a wide range of δ11B composition, from -9.3 to +44 ‰
(Aggarwal et al., 2000; Aggarwal et al., 1992; Barth, 1993; Leeman et al., 1990;
Musashi et al., 1988; Palmer and Sturchio, 1990; Vengosh et al., 1994b). The δ11B
composition of thermal water is mainly controlled by: 1) host-rock type, 2) fluid mixing,
3) B isotope fractionation and 4) steam phase separation. The last factor generally
only enriches the δ11B of up to 4 ‰ and thus is considered insignificant (Kanzaki et
al., 1979; Leeman et al., 1992; Nomura et al., 1982; Spivack et al., 1990; Yuan et al.,
2014). During thermal water-rock interaction, 11B is released into water, hence
reduces the 11B composition of the rock (Musashi et al., 1991; Palmer and Sturchio,
1990). Carbonate rocks have a wider range of δ11B composition compared to volcanic
rocks from island arc, respectively ranged from +1.5 to +26.2 ‰ (Hemming and
Hanson, 1992; Vengosh et al., 1991a) and from -2.3 to +7 ‰ (Ishikawa and
Nakamura, 1992; Palmer, 1991). B isotope fractionation at water temperatures above
65 °C was reported insignificant (Aggarwal and Palmer, 1995; Palmer et al., 1987),
thus relatively conservative during thermal water ascent. Groundwater generally has
heavier δ11B compositions than thermal water, thus groundwater dilution enriche the
δ11B of thermal water (Palmer and Sturchio, 1990; Vengosh et al., 1994a). Seawater
input has been reported elevating the δ11B value of thermal water, for instances at the
Reykjanes and Svartsengi, Iceland as well as at the Izu-Bonin arc, Kusatsu-Shirane
and Kagoshima, Japan (Aggarwal and Palmer, 1995; Aggarwal et al., 2000; Kakihana
et al., 1987; Millot et al., 2009; Musashi et al., 1988; Nomura et al., 1982; Oi et al.,
1993). Adsorption/coprecipitation of B into minerals leads to fractionation of the light 10B into minerals, thus increases the δ11B of water (Palmer et al., 1987; Schwarcz et
al., 1969; Xiao et al., 2013). B can be adsorbed/incorporated by clay minerals and iron
oxide (Lemarchand et al., 2007; Palmer et al., 1987; Schwarcz et al., 1969; Spivack
and Edmond, 1987; Vengosh et al., 1991b), calcite (Hemming and Hanson, 1992;
Vengosh et al., 1991a) and evaporite minerals (Agyei and McMullen, 1968; McMullen
et al., 1961; Oi et al., 1989; Swihart et al., 1986; Vengosh et al., 1992).
INTRODUCTION
16
GEOLOGICAL SETTING, SAMPLING AND ANALYSES
17
II. Geological setting, sampling and analyses
II.1. Geological setting The geological features of the western part of the Indonesian archipelago was
started by the collision of the Sibumasu and Indochina–East Malaya, indicated by the
distribution of Triassic granite across the island of Sumatera (Hall, 2009). This event
pushed the continental active margin, which then ceased in the early Cretaceous due
to collision of the Australian microcontinental with the Java-Meratus subduction (Hall
et al., 2007; Smyth et al., 2008). The subduction can be detected by the presence of
metamorphic rocks, stretch along Sumatera to Central Java and turn to the north to
the Borneo island (Hall, 2009) (Fig. 2.1). In the middle Eocene the subduction
reactivated to the south due to the rapid northward movement, c.a., 6-7 cm/a, of the
Australian plate (Hall, 2009; Hamilton, 1979; Müller et al., 2000; Schellart et al., 2006;
Simandjuntak and Barber, 1996). This subduction generated a volcanic arc in the
active margin, marked by the distribution of the Tertiary volcanic belt along the
southern part of Java island (Soeria-Atmadja et al., 1994; Van Bemellen, 1949) (Fig.
2.2). In the early Miocene, the subduction ceased and resumed again in the middle
Miocene, indicated by the low volcanic activities, due to the further northward
movement of the subduction hinge (Macpherson and Hall, 1999; Macpherson and
Hall, 2002; Smyth et al., 2008). The volcanic activities increased again in the late
Miocene by the formation of the Neogene volcanic arc to the north of the Tertiary
volcanic belt (Hall, 2002; Hall et al., 2007; Macpherson and Hall, 2002; Soeria-
Atmadja et al., 1994). Later the Quaternary volcanic replaced the Neogene volcanic
arc, hence the current features on Java only have two volcanic belts, the Tertiary in
the south and the Quaternary in the middle (Hamilton, 1979; Soeria-Atmadja et al.,
1994).
The volcanisms on Java produced andesitic rocks, where the younger volcanic
(Quaternary) rocks are more alkaline (Soeria-Atmadja et al., 1994). The Quaternary
volcanism produced magmas ranging from tholeiites to high-K calc alkaline, named as
‘the normal island arc association’ (Whitford et al., 1979). The resulting magma type is
associated with the distance to the subduction trench/the depth of Benioff zone
(Wheller et al., 1987; Whitford et al., 1979). Volcanoes with deeper Benioff zones
GEOLOGICAL SETTING, SAMPLING AND ANALYSES
18
produce richer K volcanic rocks. On Java, the Muria volcano has the highest K
content due to its location in the back arc (Wheller et al., 1987; Whitford et al., 1979).
Apart from the location relative to the subduction trench, the melting materials and
fluid flux also affect the typical magma produced in a volcanic island arc. Calc-alkaline
magma is produced by a high fluid flux as the result of melting of the mantle wedge
and sediment, while K-rich magma is associated with a lower fluid flux and melting of
a deeper mantle (Cottam et al., 2010). The basement of the Sunda volcanic arc are
vary from a continental crust in the West, Mesozoic accretionary complexes in the
central to east Java and an oceanic crust on Bali to Flores (Curray et al., 1977;
Hamilton, 1979). Avraham and Emery (1973) predicted the crustal thickness on Java
ranging from 20 to 25 km and thinner to the east, reached 15 km on Flores.
Fig. 2.1. Geographic and tectonic map of the Indonesian archipelago (after Hamilton, 1979 and Simandjuntak and Barber, 1996; the plates boundaries were based on Hall, 2009).
GEO
LOG
ICAL
SET
TIN
G, S
AM
PLIN
G A
ND
AN
ALY
SES
Fig.
2.2
. G
eolo
gica
l se
tting
of
the
Java
isl
and
with
the
old
Ter
tiary
vol
cani
c be
lt, Q
uate
rnar
y vo
lcan
oes
com
plex
and
fau
lts.
(ado
pted
from
Ham
ilton
,197
9; S
iman
djun
tak
and
Barb
er,1
996;
Hof
fman
n-R
othe
et a
l., 2
001;
and
Soe
ria-A
tmad
ja e
t al.,
199
4).
ES ts.
GEOLOGICAL SETTING, SAMPLING AND ANALYSES
20
The subduction in between the early Miocene to Pliocene thrust the Tertiary
volcanic belt to the north by more than 50 km (Hall et al., 2007). The continuous
subduction has also uplifted the southern part of Java. Followed by erosion, the
process thinned the crust and exposed the Tertiary (Paleogene) volcanic belt. In
Central Java the volcanic belt has been removed by excessive erosion, hence
outcropped the Cretaceous basement (Clements et al., 2009; Hall et al., 2007). This
block is the most thrust compared to the West Java and East Java due to the
presence of a couple two major strike-slip faults, the Central Java fault and Citandui
fault. The subduction pushed the East Java and West Java blocks to the north, hence
uplifted the southern part of Central Java (Bahar and Girod, 1983; Satyana, 2007;
Situmorang et al., 1976). Apart from these faults, two other major faults are present on
Java, namely the E-W backarc-thrust of Barabis-Kendeng and the NE-SW strike-slip
fault of Cimandiri (Hoffmann-Rothe et al., 2001). These faults have been generated
since the Neogene time by compressional forces (Hall, 2002; Simandjuntak and
Barber, 1996). The Cimandiri fault is an active fault with a slip rate of about 6 to 10
mm/a (Sarsito et al., 2011; Setijadji, 2010). Besides those major faults, there are
several smaller faults, which include the E-W Lembang fault in West Java, the NE-SW
Opak fault in Central Java and the NE-SW Grindulu fault in East Java.
II.2. Sampling and analyses The sampling of thermal and cold waters was performed in two periods, July to
September 2012 and October to November 2013. The first sampling covered almost
the whole area of Java, while the second was focused on Bali with some additional
samples from Java. The water samples were taken from hot spring, cold spring,
shallow thermal wells, deep geothermal well, steam vent, crater lake, freshwater lake
and seawater (e.g., Fig. 2.3).
Temperature, pH, conductivity, ORP and alkalinity, were measured in the field,
either by probe or acid titration (HACH, 2007). The samples were filtered through a
0.45 μm nylon membrane. A part of the filtered sample was used for alkalinity
measurement and two splits for the determination of anion, cation and isotopic
compositions were stored in pre-rinsed polyethylene bottles and transported to the
University of Bremen for further analyses. The cation split was acidified to 1%
concentrated HNO3 to avoid precipitation of metals. The anions, Cl-, SO42-, NO3
- and
Br-, were analyzed by ion chromatography using an IC Plus Chromathograph
(Metrohm). The cations, Ca2+, Mg2+, Na+ and K+, and Si were determined by
GEOLOGICAL SETTING, SAMPLING AND ANALYSES
21
inductively coupled plasma-optical emission spectrometry (ICP-OES) using an Optima
7300 instrument (Perkin Elmer). Trace elements of B and As were measured by using
inductively coupled plasmamass spectrometry (ICP-MS) using an iCAP-Q instrument
(Thermo Fisher). Stable isotopes (18O and 2H) were determined on a LGR DLT-100
laser spectrometer (Los Gatos Research). The isotopes results were reported in δ per
mil (‰) relative to VSMOW with an analytical uncertainty of approximately ± 1 ‰ for
δ2H and ± 0.2 ‰ for δ18O.
The B isotope composition was analyzed using a multi-collector inductively
coupled plasma mass spectrometer (MC-ICP-MS, Neptune, Thermo Fisher Scientific)
at the Isotope Geochemistry Laboratory, National Cheng Kung University, Taiwan by
following the procedure of Wang et al. (2010). A volume of 0.5 or 1 mL sample
containing a minimum of 50 ng B was used in the measurement to ensure duplicated
analysis. Prior to measurement, the HNO3 in the samples was substituted with H2O to
minimize the memory effects. B was purified from the samples by micro-sublimation
technique at 98±0.1 °C in a thermostatic hot plate rack. The 11B data were reported in
δ per mil (‰) relative to the standard of SRM NBS 951 with an analytical uncertainty
of < 0.2 ‰.
GEOLOGICAL SETTING, SAMPLING AND ANALYSES
22
Fig. 2.3. (a) Kawah Sikidang acid crater lake, Dieng; (b) Cisolok geyser; (c) Kawah Kreta steam vent, Kamojang; (d) Dieng geothermal brines; (e) Yeh Panas hot spring, Bali; (f) Tirta Husada thermal shallow wells, Bali.
(a)
(d) (c)
(b)
(e) (f)
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
23
III. Geothermal systems on the island of Java, Indonesia
Modified from:
Budi Joko Purnomo * & Thomas Pichler *
* Geochemistry & Hydrogeology, Department of Geosciences, University of Bremen, Germany
Journal of Volcanology and Geothermal Research (2014) DOI: 10.1016/j.jvolgeores.2014.08.004
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
24
Abstract
This paper presents an overview of all known geothermal systems on the island
of Java by presenting physicochemical data for associated hot springs, cold springs
and acid crater lakes. A total of 69 locations were sampled and classified based on
their position in either a volcanic complex (volcano-hosted) or a fault zone (fault-
hosted). In particular the potential of a magmatic heat source for fault-hosted
geothermal systems was investigated. Volcano-hosted geothermal systems had
higher HCO3- concentrations and higher Mg/Na ratios than fault-hosted geothermal
systems. This geochemical difference is likely due to degassing and subsequent CO2-
water reaction in the volcano-hosted systems, which is absent in the fault-hosted
geothermal systems. The HCO3 vs. Cl and Mg/Na vs. SO4/Cl systematics indicated
that fault-hosted geothermal systems located in the active Quaternary volcanic belt
received shallow magmatic fluids, hence should be classified as volcano-hosted
geothermal systems. The heat source of fault-hosted geothermal systems located in
the old (Tertiary) volcanic belt were investigated by a combination of Li enrichment
and calculated reservoir temperatures. There a shallow magmatic heat source was
only indicated for the Cilayu and Cisolok geothermal systems. Thus, a deep seated
magma was considered to be the heat source for the fault-hosted geothermal systems
of Cikundul, Pakenjeng, Parangtritis and Pacitan.
In ten of the volcano-hosted geothermal systems, 2H and 18O isotopes
enrichments were found, but not in any of the fault-hosted geothermal systems. Stable
isotope enrichment due to evaporation was recognized in the Kawah Candradimuka
and Kawah Sileri, Kawah Hujan and Candi Gedong Songo geothermal systems. A
combination of intensive evaporation and magmatic gases input produced very heavy
stable isotopes in the hot acid crater lakes of the Kawah Kamojang, Kawah Sikidang
and Kawah Putih geothermal systems. The addition of substantial amounts of
andesitic water to the geothermal fluid was observed in the Candi Songgoriti,
Banyuasin and Pablengan geothermal systems.
Contrary to established belief fault-hosted geothermal systems on Java could be
considered a potential source for geothermal energy.
Keywords:
Java, volcano-hosted and fault-hosted geothermal systems, shallow and deep
magmatic heat sources, geochemistry, stable isotope
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
25
III.1. Introduction At least 62 geothermal fields with the potential for exploitation are present on
the island of Java (Setijadji, 2010). Following Alam et al (2010) geothermal fields can
be divided into volcano-hosted and fault-hosted geothermal systems based on their
geologic association. The former is a geothermal system related to a volcanic
complex and the latter is a geothermal system located in a fault zone. To date, seven
volcano-hosted geothermal fields were developed and five of them produced
electricity. Fault-hosted geothermal fields were not developed and are rarely explored,
due to the assumption of insufficient energy. However, considering the geology of
Java, a volcanic (magmatic) influence on the fault-hosted geothermal systems is
likely.
In other volcanic arcs around the World, fault-hosted geothermal fields which
are located close to volcanic areas indicate a heating of deep circulated meteoric
water, e.g., in the Liquiñe-Ofqui fault zone of Chile and in the Southern Apennines of
Italy (Alam et al., 2010; Italiano et al., 2010). Using a trend of B enrichment, Alam et al
(2010) suggested for the Liquiñe-Ofqui fault zone (a) heating of meteoric water in
fault-zone hosted geothermal systems and (b) condensation of volcanic steam in
volcano-hosted geothermal systems. However, the authors did not indicate the heat
source of the fault hosted geothermal system. Arehart et al. (2003) identified a
magmatic heat source for the Steamboat geothermal system (Nevada, USA), based
on trace metal and gas data. Historically this geothermal system was considered as
an extensional geothermal type with anomalous heat flow as the heat source (Wisian
et al., 1999). Anomalous heat flow in the Alpine fault, New Zealand, for example, is
considered to be caused by uplift and erosion (Allis and Shi, 1995; Shi et al., 1996).
Here physicochemical processes, fluid sources and reservoir temperature of
volcano and fault-hosted geothermal systems on Java were examined, using chemical
and isotope (2H and 18O) data. The data indicated a magmatic influence on the fault-
hosted geothermal systems, and thus a hidden energy potential for some of the fault-
hosted geothermal systems on Java.
III.2. Sampling locations Water samples were collected from July to September 2012, the end of the dry
season on Java. The samples were taken from 25 geothermal systems: (1) Cisolok,
(2) Cikundul, (3) Batu Kapur, (4) Ciater, (5) Maribaya, (6) Tampomas, (7) Patuha, (8)
Pangalengan, (9) Darajat, (10) Kamojang, (11) Cipanas, (12) Kampung Sumur, (13)
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
26
Ciawi, (14) Cilayu, (15) Pakenjeng, (16) Slamet Volcano, (17) Dieng, (18) Kalianget,
(19) Ungaran, (20) Candi Dukuh, (21) Parangtritis, (22) Lawu, (23) Pacitan, (24)
Arjuna-Welirang and (25) Segaran (Fig. 3.1). In total 70 samples were collected, 61
from hot springs, 4 from cold springs, 4 from hot crater lakes and 1 from the Indian
Ocean (Table 3.1). The locations of the 4 cold spring samples were chosen based on
their proximity to those hot springs which were sampled during this investigation.
III.3. Results The results of the field and laboratory measurements are presented in Table
3.1. Cold water springs in Java were slightly acid to slightly alkaline (pH= 6.2 to 7.8)
and conductivity ranged from 86 to 324 μS/cm. Compared to the hot spring samples,
the concentrations of Ca2+, Mg2+, Na+, K+ and Cl- of the cold spring waters were low (≤
31 mg/L). These cold spring waters had HCO3- and SO4
2- contents of 19.5 to 115.9
mg/L and 2.7 to 40.6 mg/L, respectively.
The volcano-hosted hot springs had a larger variety of temperature, pH,
conductivity, major anions (HCO3-, SO4
2-, and Cl-) and two major cations (Na+ and
Mg2+), but relatively a similar range of K+ and a smaller range of Ca2+, compared to
the fault-hosted hot springs. The temperatures of the volcano-hosted hot springs
ranged from 22 to 95 °C and those of the fault-hosted hot springs ranged from 47 to
102 °C. The volcano-hosted hot springs were very acid to slightly alkaline (pH= ~ 1 to
8.4), while of the fault-hosted hot springs were slightly acid to slightly alkaline (pH= 5
to 8.1). The conductivity of the volcano-hosted hot springs varied from 86 to 14600
μS/cm, compared to 1500 to 17340 μS/cm of the fault-hosted hot springs. The
concentration of HCO3- in the volcano-hosted hot springs ranged from below detection
to 1634.8 mg/L, SO42- ranged from below detection to 3005.5 mg/L, and Cl- ranged
from 6.9 to 8084 mg/L; and those of the fault-hosted hot springs had HCO3-
concentration ranged from 22 to 1085.8 mg/L, SO42- ranged from below detection to
1284.5 mg/L, and Cl- ranged from 122.1 to 6184.5 mg/L. The concentration of Mg2+ in
the volcano-hosted hot springs ranged from 2.6 to 211.9 mg/L, Na+ ranged from 2.2 to
2979 mg/L, K+ ranged from 1.4 to 119.8 mg/L, and Ca2+ ranged from 4.9 to 510.7
mg/L; while the concentration of Mg2+ in the fault-hosted hot springs ranged from
below detection to 97.7 mg/L, Na+ ranged from 115.8 to 1797.4 mg/L, K+ ranged from
below detection to 94.2 mg/L, and Ca2+ ranged from 32.8 to 2047.6 mg/L.
Fig.
3.1
. Th
e di
strib
utio
n of
sam
pled
geo
ther
mal
sys
tem
s on
Jav
a, i
.e.,
(1)
Cis
olok
, (2
) C
ikun
dul,
(3)
Batu
Kap
ur,
(4)
Cia
ter,
(5)
Mar
ibay
a, (6
) Tam
pom
as, (
7) P
atuh
a, (8
) Pan
gale
ngan
, (9)
Dar
ajat
, (10
) Kam
ojan
g, (1
1) C
ipan
as, (
12) K
ampu
ng S
umur
, (13
) Cia
wi,
(14)
Cila
yu, (
15) P
aken
jeng
, (16
) Sla
met
Vol
cano
, (17
) Die
ng, (
18) K
alia
nget
, (19
) Ung
aran
, (20
) Can
di D
ukuh
, (21
) Par
angt
ritis
, (22
)
Law
u, (2
3) P
acita
n, (2
4) A
rjuna
-Wel
irang
and
(25)
Seg
aran
. Geo
logi
cal s
truct
ures
and
vol
cani
c be
lts w
ere
base
d on
Ham
ilton
(197
9),
Sim
andj
unta
k an
d Ba
rber
(199
6), H
offm
ann-
Rot
he e
t al.
(200
1) a
nd S
oeria
-Atm
adja
et a
l. (1
994)
.
(5)
awi,
(22)
79),
Tabl
e 3.
1. S
ampl
ing
loca
tions
, phy
sico
chem
ical
and
sta
ble
isot
ope
com
posi
tions
of c
old
sprin
gs, h
ot s
prin
gs a
nd h
ot a
cid
crat
er la
kes
on J
ava.
Sam
ple
Loca
tion
Geo
. Te
mp.
pH
Ec
TD
S C
a M
g N
a K
C
l H
CO
3 SO
4 N
O3
Br
Si
B
Li
As
18O
2 H
ID
Ty
pe
(°C
) (u
S/c
m)
(mg/
L)
(μg/
L)
(‰)
Hot
Spr
ings
J1
Pan
cura
n 3
(Sla
met
vol
c.)
V
46.3
6.
2 40
70
3995
18
9.6
204.
2 35
8.0
75.4
73
2.5
652.
7 59
9.6
<dl
<dl
85.1
4.
01
19.6
12
.1
-8.6
-6
1.9
J2
Pan
cura
n 7
(Sla
met
vol
c.)
V
52.1
6.
9 42
80
3200
20
1.2
209.
2 37
1.5
75.3
77
7.3
722.
2 61
4.7
<dl
<dl
89.6
4.
33
76.8
12
.7
-9.1
-6
0.6
J3
Cia
wi 1
V
43
.2
6.5
1500
11
23
73.0
61
.3
169.
1 43
.0
152.
7 84
4.2
<dl
<dl
<dl
86.7
5.
90
478.
5 17
.8
-6.4
-3
6.4
J4
Cia
wi 2
V
53
.4
6.7
1860
13
00
85.8
72
.5
197.
9 48
.4
165.
3 97
6.0
<dl
<dl
<dl
91.3
6.
84
534.
4 18
.3
-6.2
-3
8.9
J6
Cie
ngan
g (C
ipan
as)
V
46.2
6.
2 15
25
1048
72
.6
90.7
12
1.9
25.0
11
3.6
362.
3 45
3.4
<dl
<dl
67.2
2.
16
96.7
5.
5 -7
.4
-48.
4 J7
C
ipan
as In
dah
(Cip
anas
) V
48
.3
6.3
1632
11
32
68.7
10
7.0
143.
8 28
.6
119.
0 38
3.1
486.
4 <d
l <d
l 68
.6
2.43
12
3.4
6.6
-7.5
-4
9.8
J8
Tirta
gang
ga (C
ipan
as)
V
49.3
6.
4 16
55
1170
80
.4
100.
3 13
8.2
27.2
11
9.0
397.
7 51
3.8
<dl
<dl
70.2
2.
51
96.3
4.
7 -7
.8
-49.
7 J1
0 K
awah
Huj
an (K
amoj
ang)
V
95
.4
4.9
618
409
25.1
5.
9 19
.5
6.9
7.2
22.0
13
3.4
0.8
<dl
83.9
4.
63
0.4
3.1
-1.3
-2
1.4
J11
Tirta
husa
da (P
acita
n)
F 51
.3
5.0
4262
30
55
417.
2 <d
l 20
0.2
<dl
308.
2 23
.2
1127
.9
<dl
<dl
16.4
0.
42
23.5
30
.5
-5.7
-3
9.5
J12
Tina
tar (
Pac
itan)
F
37.3
6.
9 28
62
2115
48
4.3
2.0
185.
3 <d
l 30
8.2
22.0
12
84.5
<d
l <d
l 16
.6
0.42
26
.6
23.2
-6
.2
-35.
8 J1
3 P
adus
an 1
(Arju
na-W
elira
ng v
olc.
) V
48
.3
6.5
2574
18
82
142.
4 10
8.5
239.
9 56
.8
246.
2 11
04.1
17
1.7
<dl
<dl
74.9
4.
97
166.
9 15
.1
-9.3
-6
2.4
J14
Pad
usan
2 (A
rjuna
-Wel
irang
vol
c.)
V
45.7
6.
3 23
77
1730
11
9.5
86.8
19
2.9
46.2
20
6.3
1000
.4
140.
4 3.
9 <d
l 65
.0
4.14
13
0.8
14.0
-9
.3
-59.
3 J1
6 C
anga
r 1 (A
rjuna
-Wel
irang
vol
c.)
V
46.1
6.
8 13
36
918
72.0
87
.3
116.
4 31
.0
48.1
69
5.4
82.6
14
.2
<dl
49.9
2.
62
1.6
2.6
-8.5
-6
0.9
J17
Can
gar 2
(Arju
na-W
elira
ng v
olc.
) V
42
.3
6.7
887
599
56.0
43
.9
78.2
21
.0
38.2
59
7.8
61.9
15
.8
<dl
48.8
2.
00
1.3
1.8
-9.1
-5
8.4
J18
Son
ggor
iti 1
(Arju
na-W
elira
ng v
olc.
) V
46
.4
6.3
5470
41
78
170.
3 12
4.8
765.
2 52
.9
1303
.5
1378
.6
<dl
<dl
<dl
84.9
50
.56
1710
.4
4448
.8
-5.9
-4
6.2
J19
Son
ggor
iti 2
(Arju
na-W
elira
ng v
olc.
) V
28
.4
7.0
5549
43
42
172.
2 13
0.9
797.
5 55
.6
1366
.8
1134
.6
<dl
<dl
<dl
86.2
51
.84
1784
.4
1423
.1
-4.8
-4
4.1
J20
Son
ggor
iti 3
(Arju
na-W
elira
ng v
olc.
) V
41
.5
6.3
4658
35
14
145.
0 10
5.3
635.
6 44
.7
1084
.4
1146
.8
<dl
<dl
<dl
80.0
42
.04
1437
.2
777.
6 -6
.1
-45.
3 J2
1 S
egar
an 1
(Lam
onga
n vo
lc.)
V
44.9
6.
5 39
07
2892
10
0.7
211.
9 40
5.3
79.3
55
0.5
1625
.0
<dl
<dl
<dl
73.9
21
.37
619.
2 46
.6
-5.0
-2
9.9
J22
Seg
aran
2 (L
amon
gan
volc
.) V
22
.3
6.3
3504
25
82
91.9
18
6.8
362.
7 71
.8
497.
9 14
57.9
<d
l 3.
9 <d
l 70
.5
18.7
4 54
9.0
50.8
-5
.0
-30.
2 J2
3 C
umpl
eng
(Law
u vo
lc.)
V
34.3
6.
2 23
01
1678
73
.3
33.0
39
4.5
12.2
30
0.0
971.
1 7.
4 <d
l <d
l 52
.5
4.08
68
2.8
21.9
-6
.4
-42.
9 J2
4 B
anyu
asin
(Law
u vo
lc.)
V
38.4
6.
1 13
800
1204
0 51
0.7
146.
2 29
79.0
11
9.8
5948
.7
835.
7 25
6.4
<dl
13.0
42
.9
93.2
3 11
060.
5 95
14.8
-4
.0
-42.
4 J2
6 P
able
ngan
(Law
u vo
lc.)
V
36.4
6.
3 14
600
1275
0 13
9.5
46.4
29
67.1
81
.4
4382
.4
1634
.8
<dl
<dl
11.3
49
.6
55.6
0 34
48.9
22
9.5
-3.4
-3
1.7
J27
Nge
rak
(Law
u vo
lc.)
V
34.4
6.
2 27
56
2032
19
6.0
84.1
19
1.8
39.9
66
8.9
151.
3 24
8.3
5.1
<dl
52.5
5.
22
108.
1 6.
8 -8
.9
-59.
4 J2
8 K
ondo
(Law
u vo
lc.)
V
33.2
6.
4 48
60
3745
15
9.4
46.7
84
1.8
22.4
90
4.9
878.
4 51
1.9
<dl
<dl
30.1
17
.37
1127
.6
49.5
-5
.4
-39.
6 J2
9 B
ayan
an (L
awu
volc
.) V
39
.8
6.8
2330
16
88
103.
6 69
.9
289.
8 29
.1
323.
8 98
8.2
<dl
<dl
<dl
69.4
3.
34
297.
6 0.
9 -6
.1
-36.
5 J3
0 N
gunu
t (La
wu
volc
.) V
41
.0
6.6
2451
17
84
98.2
77
.7
316.
6 25
.4
333.
1 10
30.9
<d
l <d
l <d
l 73
.1
3.47
35
9.6
0.9
-6.3
-3
7.0
J31
Can
di D
ukuh
(Am
bara
wa)
V
35
.9
7.2
1168
80
6 46
.0
38.0
16
2.9
14.0
17
1.3
377.
0 <d
l <d
l <d
l 37
.6
3.80
74
.6
35.8
-6
.5
-44.
6 J3
2 C
andi
Son
go (U
ngar
an v
olc.
) V
48
.5
3.0
870
582
24.2
10
.0
18.9
9.
3 14
.0
<dl
340.
6 1.
0 <d
l 83
.4
0.03
6.
1 3.
8 -4
.8
-36.
3 J3
3 G
ucci
(Sla
met
vol
c.)
V
40.5
6.
3 69
4 46
4 35
.0
41.6
57
.5
24.0
33
.5
319.
6 54
.2
8.9
<dl
59.9
2.
66
19.6
5.
3 -8
.7
-60.
4 J3
4 P
enga
siha
n (S
lam
et v
olc.
) V
53
.3
7.5
1206
81
9 46
.0
56.5
14
4.5
41.0
52
.6
556.
3 92
.0
<dl
<dl
75.4
7.
15
76.8
7.
9 -8
.3
-63.
8 J3
5 M
arib
aya
F 46
.5
6.2
2345
16
60
129.
0 97
.7
115.
8 26
.4
122.
1 10
17.5
<d
l <d
l <d
l 82
.5
1.94
11
6.5
1.2
-7.5
-5
0.1
J36
Cia
ter
V
46.6
2.
0 63
11
4850
88
.6
30.0
58
.7
41.4
82
2.5
<dl
659.
2 6.
0 <d
l 78
.7
2.07
37
.6
68.3
-6
.9
-44.
6 J3
7 Ba
tu K
apur
1
F 56
.3
6.7
2154
15
30
50.6
46
.2
310.
4 58
.5
280.
4 71
7.4
75.4
<d
l <d
l 87
.9
2.74
53
9.7
1.3
-7.1
-4
1.4
J38
Batu
Kap
ur 2
F
40.9
6.
4 24
86
1802
10
2.1
96.6
27
6.9
39.9
31
2.7
1085
.8
<dl
<dl
<dl
75.1
3.
12
703.
4 2.
5 -6
.4
-39.
3 J3
9 C
ibol
ang
(Pan
gale
ngan
) V
68
.9
7.1
988
650
77.0
43
.0
71.3
28
.0
24.0
21
9.6
236.
0 <d
l <d
l 95
.5
6.48
63
.8
14.2
-6
.5
-47.
3 J4
0 S
ukar
atu
(Pan
gale
ngan
) V
39
.8
6.4
1005
68
2 64
.0
53.2
97
.8
14.0
18
.6
475.
8 10
2.6
<dl
<dl
70.1
2.
12
61.5
6.
2 -8
.1
-56.
1 J4
1 K
erta
man
ah (P
anga
leng
an)
V
54.1
6.
4 32
08
553
54.0
35
.7
94.4
26
.0
17.6
43
3.1
91.8
<d
l <d
l 92
.5
1.06
56
.1
35.8
-8
.1
-53.
3
2 H‰
) -61.
9-6
0.6
-36.
4-3
8.9
-48.
4-4
9.8
-49.
7-2
1.4
-39.
5-3
5.8
-62.
4-5
9.3
-60.
9-5
8.4
-46.
2-4
4.1
-45.
3-2
9.9
-30.
2-4
2.9
-42.
4-3
1.7
-59.
4-3
9.6
-36.
5-3
7.0
-44.
6-3
6.3
-60.
4-6
3.8
-50.
1-4
4.6
-41.
4-3
9.3
-47.
3-5
6.1
-53.
3
Ta
ble
3.1.
(con
tinue
d)
Sam
ple
Loca
tion
Geo
. Te
mp.
pH
Ec
TD
S C
a M
g N
a K
C
l H
CO
3 SO
4 N
O3
Br
Si
B
Li
As
18O
2 H
ID
Ty
pe
(°C
) (u
S/c
m)
(mg/
L)
(μg/
L)
(‰)
Hot
Spr
ings
J4
2 P
aken
jeng
F
59.9
7.
4 19
60
1367
22
5.5
<dl
224.
3 <d
l 12
6.0
40.3
94
0.2
<dl
<dl
27.1
7.
21
92.8
68
5.4
-6.5
-3
5.6
J43
Pak
enje
ng
F 43
.1
7.5
2048
14
61
215.
6 <d
l 25
6.9
<dl
131.
7 42
.7
960.
0 <d
l <d
l 26
.6
7.68
72
.4
699.
7 -6
.2
-35.
8 J4
4 C
ilayu
F
70.3
8.
1 58
92
4435
68
.3
12.9
11
01.5
66
.1
1387
.2
372.
1 40
8.1
<dl
4.9
79.4
58
.22
2233
.0
3521
.5
-5.3
-4
0.8
J45
Cila
yu
F 45
.1
7.9
1014
0 83
16
227.
0 9.
9 17
97.4
94
.2
3210
.5
289.
1 15
6.6
<dl
11.1
82
.8
47.5
7 17
80.6
27
79.3
-5
.3
-33.
9 J4
6 K
alia
nget
V
38
.9
6.6
2611
19
27
118.
2 15
0.7
190.
1 45
.3
399.
1 84
4.2
150.
8 <d
l <d
l 61
.7
3.52
20
3.4
135.
9 -9
.1
-58.
4 J4
7 K
alia
nget
V
40
.0
6.5
2657
19
54
128.
3 15
6.7
197.
5 49
.3
424.
8 73
2.0
163.
6 <d
l <d
l 66
.2
3.73
20
8.3
189.
3 -9
.0
-59.
5 J4
8 C
ikun
dul
F 50
.5
7.8
1500
10
26
32.8
1.
1 26
4.3
5.4
180.
2 61
.0
374.
2 <d
l <d
l 37
.6
10.7
2 94
.2
38.8
-5
.7
-39.
4 J4
9 C
isol
ok
F 10
2.0
8.1
1705
11
80
41.2
3.
0 28
5.8
10.3
30
5.6
129.
3 23
5.5
<dl
<dl
66.3
3.
58
290.
1 10
4.0
-5.9
-3
3.0
J50
Cis
olok
F
100.
0 8.
0 16
12
1078
52
.3
3.3
257.
6 8.
8 27
7.0
161.
0 22
2.7
<dl
<dl
58.9
3.
20
251.
9 96
.1
-6.0
-3
5.3
J52
Pat
uha
V
64.3
8.
4 71
5 46
8 41
.6
16.7
71
.4
16.5
35
.2
300.
1 46
.9
<dl
<dl
80.7
1.
29
44.9
43
.1
-8.9
-5
6.6
J53
Tam
pom
as
V
51.4
6.
9 19
12
1349
73
.2
50.8
24
1.9
22.8
28
0.2
705.
2 1.
6 <d
l <d
l 90
.8
4.95
49
0.5
0.7
-6.2
-4
2.8
J54
Tam
pom
as
V
48.5
7.
1 34
62
2526
83
.4
50.6
54
2.2
30.2
75
7.2
732.
0 <d
l <d
l 2.
8 82
.7
5.28
15
23.0
2.
1 -6
.6
-41.
2 J5
5 D
araj
at
V
60.0
2.
8 95
2 64
3 13
.7
5.3
7.9
2.9
13.3
<d
l 25
4.1
1.0
<dl
78.8
6.
97
2.9
87.2
-7
.9
-50.
6 J5
6 D
araj
at
V
54.0
3.
8 34
4 22
2 24
.0
4.3
10.9
5.
0 6.
9 <d
l 11
7.6
0.6
<dl
49.7
1.
11
5.8
17.1
-8
.7
-51.
9 J5
7 K
ampu
ng S
umur
V
35
.0
7.6
592
399
19.0
18
.1
85.9
14
.0
52.0
22
9.4
13.9
0.
9 <d
l 43
.0
0.95
33
.1
10.2
-7
.5
-47.
2 J5
8 Pa
rang
tritis
F
39.2
7.
6 17
340
1543
0 20
47.6
8.
7 16
40.0
21
.4
6184
.5
43.9
47
7.0
<dl
18.1
26
.2
9.51
28
2.3
15.6
-4
.3
-24.
2 J6
0 D
ieng
V
57
.4
6.3
1104
76
2 60
.0
19.5
11
2.4
27.0
77
.7
266.
0 20
2.9
0.8
<dl
49.5
6.
67
13.8
6.
1 -4
.3
-47.
4 J6
1 D
ieng
V
54
.0
6.2
1550
10
82
130.
4 40
.2
104.
1 59
.1
330.
6 18
3.0
67.8
<d
l 1.
0 10
3.5
6.41
52
.8
2.1
-8.0
-5
7.2
J62
Die
ng
V
70.1
6.
8 34
3 21
6 23
.8
11.2
16
.6
20.3
9.
0 12
4.4
16.6
<d
l <d
l 97
.9
0.09
2.
5 91
.1
-8.6
-5
7.6
J63
Die
ng
V
56.1
6.
7 74
5 48
7 41
.9
25.1
76
.0
26.5
27
.3
270.
8 11
9.7
<dl
<dl
91.6
4.
55
22.6
2.
3 -7
.3
-55.
6 J6
4 D
ieng
V
60
.0
7.3
744
487
41.7
14
.3
95.5
32
.2
21.5
32
9.4
82.8
0.
9 <d
l 81
.4
2.08
45
.0
6.6
-8.0
-5
7.1
J65
Die
ng
V
81.0
7.
3 15
75
1046
16
.3
6.6
4.0
2.2
14.7
10
4.9
298.
5 1.
3 <d
l 13
.2
0.54
6.
7 1.
2 0.
6 -2
1.0
J66
Die
ng
V
31.6
5.
9 42
7 28
3 35
.6
21.6
15
.7
11.2
15
.7
201.
3 22
.4
0.5
<dl
55.8
0.
27
11.2
3.
8 -8
.2
-54.
8 J6
8 D
ieng
V
23
.4
2.5
1748
12
60
4.9
2.6
2.2
1.4
13.1
<d
l 43
4.0
0.9
<dl
45.4
0.
15
3.6
2.4
-8.6
-5
4.3
J69
Die
ng
V
26.6
6.
0 40
0 26
7 35
.2
24.3
10
.9
8.1
16.8
17
9.3
41.2
<d
l <d
l 52
.0
0.17
4.
4 0.
3 -9
.3
-62.
3 H
ot C
rate
r Lak
es
J9
Kam
ojan
g V
40
.0
2.9
1185
82
6 63
.0
23.5
15
.9
23.0
12
.8
<dl
406.
8 <d
l <d
l 16
8.8
1.39
13
.3
1.3
7.7
-4.1
J5
1 P
atuh
a V
32
.9
1.0
86
115
75.0
32
.1
34.2
33
.0
8084
.2
<dl
3005
.5
0.7
20.5
12
2.4
94.4
0 41
.7
236.
5 7.
9 -4
.0
J59
Die
ng
V
86.8
2.
5 25
69
1851
18
.3
14.5
11
.4
5.9
14.4
<d
l 90
0.6
<dl
<dl
161.
5 73
.29
12.6
1.
4 7.
5 -7
.6
Col
d Sp
rings
J5
C
iaw
i -
25.2
6.
5 16
9 10
8.1
15.0
9.
5 2.
9 <d
l 9.
1 97
.6
2.7
0.9
<dl
36.0
0.
01
2.7
0.9
-6.1
-4
2.0
J15
Arju
na-W
elira
ng
- 22
.6
6.3
324
215.
0 31
.0
13.6
16
.0
5.0
22.0
11
5.9
30.4
8.
2 <d
l 30
.7
0.27
2.
6 2.
9 -7
.6
-52.
5 J2
5 La
wu
- 21
.9
6.2
86
55.3
5.
0 1.
6 9.
5 <d
l 9.
5 19
.5
12.7
2.
2 <d
l 20
.5
0.10
22
.0
73.2
-7
.9
-52.
8 J6
7 D
ieng
-
18.1
7.
6 28
6 19
0.4
28.8
6.
6 13
.5
4.8
15.6
15
.9
40.6
48
.4
<dl
21.6
0.
97
1.4
0.9
-8.8
-5
5.6
Seaw
ater
SW
In
dian
Oce
an
- nm
nm
nm
nm
36
9.7
1372
.8
1068
3 36
0 19
191
151
2183
<d
l 60
.3
nm
9.96
11
.6
nm
nm
nm
*V=
volc
ano-
host
ed, F
= fa
ult-h
oste
d, v
olc.
= vo
lcan
o, <
dl=
belo
w d
etec
tion
limit,
nm
= no
t mea
sure
d
2 H) -3
5.6
-35.
8-4
0.8
-33.
9-5
8.4
-59.
5-3
9.4
-33.
0-3
5.3
-56.
6-4
2.8
-41.
2-5
0.6
-51.
9-4
7.2
-24.
2-4
7.4
-57.
2-5
7.6
-55.
6-5
7.1
-21.
0-5
4.8
-54.
3-6
2.3
-4.1
-4.0
-7.6
-42.
0-5
2.5
-52.
8-5
5.6
nm
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
30
Both the volcano- and fault-hosted hot springs were characterized by relatively
large variation of B, Li and As concentrations. The B concentration of the volcano-
hosted systems ranged from 0.03 to 94.4 mg/L, Li ranged from 0.4 μg/L to 11.06 mg/L
and As ranged from 0.3 μg/L to 9.5 mg/L. In the fault-hosted systems, the B
concentrations ranged from 0.42 to 58.2 mg/L, Li ranged from 23.5 μg/L to 2.23 mg/L
and As ranged from 1.2 μg/L to 3.5 mg/L. Most of the hot springs with high B, Li and
As concentrations were chloride water. This phenomenon is common in geothermal
systems, because neutral chloride waters ascend directly from the reservoir and thus
are generally enriched in selected trace elements, i.e., the geothermal suite of
elements (Goff and Janik, 2000; Nicholson, 1993; White et al., 1971).
In addition to physicochemical parameters, stable isotopes of 2H and 18O were
determined. The stable isotope composition of hot spring waters from the volcano-
hosted systems had a larger variation than those from the fault-hosted systems. The 18O isotope composition of the cold springs ranged from -8.8 to -6.1 ‰, the volcano-
hosted hot springs waters ranged from -9.3 to 7.9 ‰ and the fault-hosted hot springs
waters ranged from -7.5 to -4.3 ‰ (Table 3.1). The 2H isotope composition of the cold
springs ranged from -55.6 to -42.0 ‰, the volcano-hosted hot springs waters ranged
from -62.3 to -4.1 ‰ and the fault-hosted hot springs waters ranged from -50.1 to -4.2
‰ (Table 3.1).
III.4. Discussion III.4.1. General considerations about geothermal systems on Java
As mentioned above, geothermal systems on Java were classified into volcano-
hosted and fault-hosted. Based on this classification, from a total of 25 sampled
geothermal systems, 8 were considered fault-hosted (i.e., Pacitan, Maribaya, Batu
Kapur, Pakenjeng, Cilayu, Cikundul, Cisolok, and Parangtritis) and 17 were
considered volcano-hosted (i.e., Segaran, Arjuna-Welirang Volcano, Lawu Volcano,
Ungaran Volcano, Candi Dukuh, Dieng, Kalianget, Slamet Volcano, Ciawi, Kampung
Sumur, Tampomas, Cipanas, Ciater, Darajat, Kamojang, Pangalengan, and Patuha)
(Fig. 3.1). All of the volcano-hosted geothermal systems were in the Quaternary
volcanic belt, while most of the fault-hosted geothermal systems were in the Tertiary
volcanic belt (Fig. 3.1).
Several of those fault-hosted geothermal systems are located in major fault
zones, e.g. the Cisolok and Cikundul geothermal systems in the Cimandiri fault, the
Maribaya geothermal system in the Lembang fault, the Parangtritis geothermal
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
31
system in the Opak fault, and the Pacitan geothermal system in the Grindulu fault.
The Batu Kapur, the Pakenjeng and the Cilayu geothermal systems are associated
with minor faults (Fig. 3.1). Contrast with the other fault-hosted geothermal systems,
the Maribaya and the Batu Kapur were located in the active Quarternary volcanic belt,
thus probably are heated by volcanic activity. According to the geologic maps of Java
(Alzwar et al., 1992a; Effendi et al., 1998a; Samodra et al., 1992; Silitonga, 1973a;
Sujatmiko and Santosa, 1992a), the Pacitan, the Pakenjeng, the Cilayu, the Cisolok
and the Parangtritis geothermal systems are situated close to the zones of the
Tertiary intrusive rocks. The position of a fault-hosted geothermal field nearby an
intrusive rock could be an indication of a magmatic heat source, an assumption that
was further investigated using geochemical tools.
The origin and physicochemical history of hydrothermal fluids can be explored in
a Cl, SO4 and HCO3 ternary diagram (Chang, 1984; Giggenbach, 1991; Giggenbach,
1997; Nicholson, 1993). Based on their position in the diagram, hydrothermal waters
can be divided into neutral chloride, acid sulfate and bicarbonate waters, but mixtures
between the individual groups are common. On Java, bicarbonate was the dominant
water type for the volcano-hosted hot springs, followed by acid sulfate and neutral
chloride waters, while for the fault-hosted hot springs the water types were distributed
more or less evenly (Fig. 3.2). The occurrence of different water types in a given
hydrothermal system is common, indicating the different physicochemical processes,
such as, phase separation and mixing in the shallow subsurface (Ellis and Mahon,
1977; Giggenbach, 1997; Hedenquist, 1990; Henley and Ellis, 1983; McCarthy et al.,
2005). Although the bicarbonate water type seems to be more abundant in the
volcano-hosted hydrothermal systems, a definite difference between the volcano- and
fault-hosted systems is difficult to be assessed in a ternary diagram alone. Thus,
following the procedure of Valentino and Stanzione (2003), the hot waters from Java
were plotted in HCO3 vs. Cl and Mg/Na vs. SO4/Cl diagrams (Figs. 3.3 and 3.4).
These diagrams show that the volcano-hosted thermal waters have a higher HCO3-
content and a higher Mg2+/Na+ ratio. This observation could result due to magmatic
degassing and thus addition of CO2 to the volcano-hosted hot springs, which is likely
absent or minor in the fault-hosted hot springs. The reaction between H2O and CO2
increases acidity and thus intensifies water-rock interaction (Giggenbach, 1984;
Giggenbach, 1988; White, 1957). Acid conditions and low temperature, due to slow
upward migration or a long flow path of the ascending thermal waters, increases the
solubility of Mg2+ (Allen and Day, 1927), hence producing Mg-rich waters.
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
32
Fig. 3.2. SO4-HCO3-Cl ternary diagram of cold and hot springs. Most of the volcano-hosted hot springs were of the HCO3
- water type, whereas the fault-hosted hot springs were distributed evenly between the SO42-, HCO3
- and Cl- types. Open circle = cold spring, gray filled triangle = volcano-hosted hot spring and open square = fault-hosted hot spring.
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
33
Fig. 3.3. HCO3 vs. Cl (in mg/L basis) diagram of cold and hot springs for: (A) and (B) formed in the margin of the ‘primary neutralization’, where (A) are closer than (B), (C) thermal waters from shallow depth, thus undergone major dilution by groundwater, (D) volcanic H2S gas oxidized by O2-rich groundwater, (E) and (F) fault-hosted hot springs, where (E) were less diluted by groundwater than the (F). (E) is likely influenced by seawater input. (Symbols are as in Fig. 3.2).
Fig. 3.4. Mg/Na vs. SO4/Cl (in meq/L basis) diagram of cold and hot springs. The interpretation of groups A, B, C, D, E and F are similar to those in Fig. 3.4. (Symbols are as in Fig. 3.2).
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
34
Four groups of the volcano-hosted thermal waters are shown in the HCO3 vs. Cl
and the Mg/Na vs. SO4/Cl diagrams, i.e., A, B, C and D (Fig. 3.3 and 3.4). The group
A samples had the highest HCO3- and Cl- concentrations, but the lowest SO4/Cl ratios.
Conversely, the group D samples had the lowest HCO3- and Cl- concentrations but the
highest SO4/Cl ratios. Meanwhile, the group B and C had HCO3- and Cl-
concentrations and SO4/Cl ratios in between of the group A and D. However, the
group C was split from the group B due to its lower HCO3- and Cl- contents. The
groups A and B are thought to have formed at the margin of the ‘primary
neutralization’ zone (Giggenbach, 1988). There separation of CO2 and its reaction
with groundwater produces HCO3-rich thermal waters, while the Cl- content remains
high due to the lesser dilution by groundwater. The group A samples originated closer
to the ‘primary neutralization’ zone than the group B samples and thus had a higher
Cl- concentration. In the Mg/Na vs. SO4/Cl diagram, two acid thermal waters, J51
(Kawah Putih) and J36 (Ciater), are exceptions in group B. The moderate SO42-/Cl-
ratios in these two samples were likely caused by H2S and HCl addition to shallow
groundwater (Delmelle and Bernard, 1994; Delmelle et al., 2000; Giggenbach, 1988;
Sriwana et al., 2000). The resulting low pH enhances water-rock interaction, thus
causing the high Mg2+/Na+ ratio in those two samples. The group C samples were
considered thermal waters formed in the shallow subsurface and thus were diluted by
groundwater, which lowered their HCO3- and Cl- contents. The group D samples were
interpreted to be thermal waters influenced by primary H2S-rich magmatic vapor in the
shallow subsurface, hence producing acid sulfate waters. In this group, J65 (Dieng)
was an exception (Fig. 3.4), because the hot spring had a neutral (pH= 7) and was Cl-
poor (14.7 mg/L). This thermal water is likely a mixture of HCO3-rich and SO4-rich
water, which can occur in low relief liquid-dominated geothermal systems (Kuhn,
2004). Meanwhile, the thermal waters of the fault-hosted geothermal systems were
divided into two groups, E and F, where the former had a higher Cl- content than the
latter. Chloride is a conservative element and thus the Cl- variation in the fault-hosted
thermal waters could be caused by either varying degrees of mixing with shallow
groundwater or reflect the initial Cl- content of the hydrothermal fluid. Based on that
assumption, group E thermal waters should have undergone less mixing with shallow
groundwater than group F. This is also indicated in the Na-K-Mg ternary diagram
(Giggenbach, 1988), where the group E thermal waters (J45 and J58) plot closer to
equilibrium than the group F thermal waters (Fig. 3.6). The high Cl- concentration in
sample J58 from the Parangtritis hot spring is unusual, but can be explained by
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
35
seawater addition to the geothermal system, because of its proximity to the Indian
Ocean. Most of the acid sulfate waters had HCO3- concentrations, which were below
detection and thus less samples could be plotted as group D in the HCO3 vs. Cl
diagram than in the Mg/Na vs. SO4/Cl diagram. Conversely, many neutral chloride
waters were below the detection limit of SO42- and therefore more thermal waters
plotted as group A in the HCO3 vs. Cl diagram than in the Mg/Na vs. SO4/Cl diagram.
Several anomalies and discrepancies were found in both the HCO3 vs. Cl and
Mg/Na vs. SO4/Cl diagrams. In the HCO3 vs. Cl diagram, the fault-hosted hot springs
of Maribaya (J35) and Batu Kapur (J37 and J38) plot within the group B of the
volcano-hosted thermal waters, whereas the volcano-hosted hot springs of J27
(Lawu) and J61 (Dieng) plot as a group of fault-hosted thermal waters (Fig. 3.3). As
mentioned previously, the HCO3- abundance in the Maribaya and Batu Kapur hot
springs was likely caused by addition of magmatic CO2, due to their location within the
active Quaternary volcanic belt. Meanwhile, the lower HCO3- concentrations in
samples J27 (Lawu) and J61 (Dieng) were likely the result of carbonate mineral
precipitation. That sample J28 plots below group B in the Mg/Na vs. SO4/Cl diagram
should be due to the formation of clay minerals and the associated depletion of Mg2+.
The discrepancy that samples J23 and J53 plot in group B in the HCO3 vs. Cl diagram
and in group A in the Mg/Na vs. SO4/Cl diagram can be attributed to the removal of
SO4 due to precipitation of sulfate minerals. The same process is the likely cause for
the location of sample J57, a dilute thermal water, in group B in the Mg/Na vs. SO4/Cl
diagram.
The Cl/B ratios of hydrothermal waters can be used to identify subsurface
processes, such as, water-rock interaction, magma degassing and seawater feeding
in a geothermal system (Arnorsson and Andresdottir, 1995; Valentino and Stanzione,
2003). The Cl/B ratios found in the samples from Java indicated three dominant
processes for the volcano-hosted thermal waters, i.e., groundwater mixing, water-rock
interaction with the andesitic host rock and phase separation. On the other hand, the
Cl/B ratios of fault-hosted hot springs were generally affected by water-rock
interaction with the andesitic host rock (Fig. 3.5).
Geothermal water in the Dieng, Kamojang and Darajat geothermal fields had
lower Cl/B ratio than the andesitic rock (Fig. 3.5), something that can be caused by
phase separation in high temperature (>300 °C) reservoirs (Truesdell et al., 1989).
This process removes B from the geothermal reservoir, thus relatively increasing the
Cl- concentration of the remaining hydrothermal fluid (Arnorsson and Andresdottir,
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
36
1995; Truesdell et al., 1989). One volcano-hosted hot spring, J36 (Ciater), and three
fault-hosted hot springs, J11 (Pacitan), J12 (Pacitan) and J58 (Parangtritis), plotted
closed to the seawater-precipitation line. The J58 (Parangtritis) sample also plotted in
the HCO3 vs. Cl and Mg/Na vs. SO4/Cl diagrams in a way which would indicate
seawater addition. However, seawater addition is not likely for the J12 and J13 hot
springs due to their low Cl- concentration. Hence, the low B/Cl ratio of these hot
springs could have been caused by B removal due to adsorption by clay minerals,
particularly illite (Harder, 1970).
Fig. 3.5. Cl vs. B diagram of cold waters and geothermal waters. The Cl/B ratio of seawater from the Indian Ocean (this research) and andesitic rock from Trompetter et al. (1999) are used. The Cl/B ratio of volcano-hosted and fault-hosted hot springs were considered to be controlled by water-rock interaction. The Dieng, Kamojang and Darajat volcano-hosted geothermal systems underwent phase separation. Though J58 (Parangtritis), J11 and J12 of Pacitan and J36 (Ciater) plotted close to seawater, but only for J58 seawater mixing was indicated. (Symbols are as in Fig. 3.2).
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
37
III.4.2. Geothermometry Solute geothermometers, such as listed in Table 3.2, can provide powerful tools
to estimate subsurface conditions. Their successful application has been extensively
discussed in the geothermal literature and relies on five basic assumptions: 1)
exclusively temperature dependent mineral-fluid reaction; 2) abundance of the mineral
and/or solute; 3) chemical equilibrium; 4) no re-equilibration; and 5) no mixing or
dilution (e.g., Nicholson, 1993). The no mixing or dilution assumption, however, can
be circumvented if their extent and/or influence on solute ratios (e.g., Na/K) are
known. Several geothermometers were applied in order to analyze the
physichochemical processes encountered by the hydrothermal fluids during their
ascent to the surface. These processes include dilution by shallow water, conductive
cooling, adiabatic cooling, mineral precipitation, adsorption/desorption, water-rock
interaction and re-equlibrium (Fournier, 1977; Kaasalainen and Stefánsson, 2012).
Different geothermometers record different equilibria and disagreement does not
immediately eliminate the use of one or the other. Careful application and evaluation
of calculated temperatures may provide important clues to the overall hydrology of the
geothermal system (Pichler et al., 1999).
Silica geothermometers, which are commonly applied to hot springs (Fournier,
1977) predicted a temperature range from 100 to 140 °C for most samples (Table
3.2). Lower temperatures were calculated for J60 (Kampung Sumur), J31 (Candi
Dukuh), and J24 and J28 (Lawu). The J60 and J31 hot springs were located at the
edge of a pool and a lake, respectively and thus, should be diluted by surface water.
The two hot springs, J24 and J28, were Cl-rich, which would preclude dilution by
surface or groundwater dilution. That lead to the conclusion that the J24 and J28 hot
springs lost silica due to the precipitation of silicate minerals during ascent, causing
the low silica geothermometer temperatures.
When applied to hot springs, silica thermometers are known to predict closer to
the discharge temperature, rather than the reservoir temperature (e.g., Pichler et al.,
1999). This inherent problem can be overcome by calculating the silica ‘parent’
concentration using the silica mixing model of Fournier (1977). After application, the
silica geothermometers predicted much higher reservoir temperatures of 258 °C for
Slamet, 188 °C for Ciawi, 180 °C for Batu Kapur and 221 °C for Pangalengan. These
temperatures were in the range of those, predicted by the Na/K and Na/K/Ca
geothermometers (see below).
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
38
Table 3.2. Calculated reservoir temperatures.
Location Samples Geothermal T Field Geothermometers (°C)
ID Types (oC) Quartz Quartz Chalcedony Quartz Na/K Na-K-Ca
(steam
loss) (parents)
Slamet
J1
V
46.3 128 125 101
258
290 217 J2 52.1 131 128 104 286 215 J33 40.5 110 110 81 380 241 J34 53.3 122 120 94 326 232
Ciawi J3 V 43.2 129 126 102 188 313 223 J4 53.4 132 128 105 308 222
Cipanas J6
V 46.2 116 115 87
nd 287 203
J7 48.3 117 116 88 284 205 J8 49.3 118 117 90 283 202
Arjuna-Welirang
J13
V
48.3 122 119 93
nd
304 219 J14 45.7 114 113 85 305 217 J16 46.1 102 102 72 318 218 J17 42.3 101 102 71 319 213 J18 46.4 128 125 101 187 169 J19 28.4 129 126 101 188 180 J20 41.5 125 122 97 189 168
Segaran J21 V 44.9 121 119 93 nd 282 222 J22 22.3 118 117 90 283 221
Lawu
J23
V
34.3 104 104 74
nd
133 128 J24 38.4 95 96 64 150 155 J26 36.4 101 102 71 127 147 J27 34.4 104 104 74 289 203 J28 33.2 79 83 48 125 126 J29 39.8 118 116 89 217 176 J30 41.0 120 118 92 199 166
Candi Dukuh J31 V 35.9 89 91 58 nd 204 165
Pangalengan J39
V 68.9 135 131 108
221 371 232
J40 39.8 118 117 90 250 180 J41 54.1 133 129 106 323 219
Kalianget J46 V 38.9 112 111 83 nd 305 216 J47 40.0 115 114 86 310 219
Patuha J52 V 32.9 125 123 98 nd 301 205
Tampomas J53 V 64.3 132 128 105 nd 212 172 J54 51.4 127 124 99 171 158
Kampung Sumur J57 V 35.0 95 96 64 nd 263 196
Dieng
J60
V
57 101 102 71
nd
306 213 J61 54 139 134 113 431 261 J62 70 136 132 109 599 296 J63 56 132 129 105 354 233 J64 60 126 123 98 349 236 J66 32 107 107 77 474 248 J69 27 104 104 74 481 242
Pacitan J11 F 51.3 56 62 24 nd nd nd J12 37.3 54 63 24 Maribaya J35 V 46.5 127 124 99 nd 299 203
Batu Kapur J37 V 56.3 130 127 103 180 278 221 J38 40.9 122 120 93 250 195
Pakenjeng J42 F 59.9 75 79 44 nd nd nd J43 43.1 74 79 43
Cilayu J44 F 70.3 125 122 97 nd 177 176 J45 45.1 127 124 99 167 167
Cikundul J48 F 50.5 89 91 58 nd 111 111
Cisolok J49 F 102.0 115 114 87 nd 143 133 J50 100.0 110 109 80 139 128
Parangtritis J58 F 39.2 74 78 42 nd 88 55 *V= volcano-hosted, F= fault-hosted, nd= not defined
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
39
Reservoir temperature calculations with the Na/K and Na/K/Ca
geothermometers were conducted according to Giggenbach (1988), by first evaluating
if equilibrium between host-rock and hydrothermal fluid was attained. Only six
samples, J24 and J26 (Lawu volcano), J44 and J45 (Cilayu), J48 (Cikundul), and J58
(Parangtritis) were in partial equilibrium, and four samples, J23 and J28 (Lawu), and
J49 and J50 (Cisolok) were close to partial equilibrium with their respective host rocks
(Fig. 3.6). The Na/K geothermometer was then applied for these ten samples and
calculated reservoir temperatures were 127 to 150 °C for Lawu, 167 to 177 °C for
Cilayu, 111 °C for Cikundul, 88 °C for Parangtritis and 139 to 143 °C for Cisolok.
Fig. 3.6. Na-K-Mg ternary diagram Giggenbach (1988) for the Java hot springs. Several fault-hosted and a few of volcano-hosted hot springs were close to or in partial equilibrium with the host-rock. (Symbols are as in Fig. 3.2).
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
40
Fournier (1989) respectively proposed the use of the Na/K geothermometer
and the Na-K-Ca geothermometer for the prediction of the highest and the lowest
reservoir temperatures in a geothermal system. Following this approach, calculated
temperatures ranged from 205 to 301 °C in Patuha and 219 to 323 °C in Pangalengan
geothermal systems. These results were in good agreement with the predicted
reservoir temperatures by direct measurement, i.e., 209 to 241 °C for Patuha (Layman
and Soemarinda, 2003) and 250 to 300 °C for Pangalengan (Abrenica et al., 2010;
Layman and Soemarinda, 2003). Reliable reservoir temperature prediction with those
two geothermometers was also indicated in the Dieng geothermal system, where
calculated temperatures ranged from 236 to 349 °C. These temperatures were
relatively similar to the predictions by Prasetio et al. (2010) (i.e, 240 to 333 °C), who
used gas geothermometers. Therefore, based on those observation, the two
geothermometers were applied to the remaining geothermal systems and calculated
temperatures in Arjuna-Welirang ranged from 217 to 305 °C, in Cipanas from 202 to
277 °C, in Segaran from 221 to 283 °C, in Kalianget from 216 to 305 °C, in Tampomas
from 172 to 212 °C and in Maribaya from 203 to 299 °C.
The K+ concentration in the Pacitan and Pakenjeng hot springs waters were
below detection limit, thus the Na/K and Na/K/Ca geothermometers could not be
applied. Silica geothermometers resulted in a maximum temperature of 63 °C for
Pacitan and 79 °C for Pakenjeng geothermal systems. Those temperatures although
likely lower than the actual reservoir temperatures indicated reservoir temperatures
below 100 °C. A compilation of all calculated reservoir temperatures on Java are
presented in Table 3.3.
In the Darajat and Kamojang geothermal systems, the only surface expressions
were acid sulfate-type hot springs. That type of hydrothermal fluid reacts extensively
with near surface rocks, hence chemical geothermometers could not be applied (e.g.
Nicholson, 1993). A reservoir temperature of 280 °C was predicted by Hadi (1997) for
the Darajat geothermal system, based on the alteration minerals observed in drill
cores and Sudarman et al. (1995) reported a shallow reservoir temperature
measurement of 232 °C for the Kamojang geothermal system.
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
41
Table 3.3. Compilation of calculated geothermal reservoir temperatures on Java.
Geothermal Geothermal T (oC) Geothermometer systems Types
Slamet M. V 258 to 380 Si parent and Na-K
Ciawi V 188 to 313 Si parent and Na-K
Cipanas V 202 to 287 Na-K and Na-K-Ca
Arjuna-Welirang M. V 217 to 305 Na-K and Na-K-Ca
Segaran V 221 to 283 Na-K and Na-K-Ca
Lawu M. V 127 to 150 Na-K
Candi Dukuh V 165 to 204 Na-K and Na-K-Ca
Pangalengan V 221 to 323 Si parent and Na-K
Kalianget V 216 to 310 Na-K and Na-K-Ca
Patuha V 205 to 301 Na-K and Na-K-Ca
Tampomas V 172 to 212 Na-K and Na-K-Ca
Kampung Sumur V 196 to 263 Na-K and Na-K-Ca
Dieng V 236 to 349 Na-K and Na-K-Ca
Pacitan V <100 Si
Maribaya F 203 to 299 Na-K and Na-K-Ca
Batu Kapur F 180 to 278 Si parent and Na-K
Pakenjeng F <100 Si
Cilayu F 125 to 177 Na-K
Cikundul F 111 Na-K
Cisolok F 139 to 143 Na-K
Parangtritis F 88 Na-K
*V= volcano-hosted, F= fault-hosted
III.4.3. The heat sources of the fault-hosted geothermal systems
The Quaternary volcanic arc could be the heat source for the fault-hosted
geothermal systems. That case was investigated by comparing the enrichment of
conservative trace elements and reservoir temperatures between the volcanic and
fault-hosted geothermal fields. A similar procedure was applied to the Steamboat
geothermal system (Arehart et al., 2003), which pointed towards a magmatic heat
source rather than just enhanced heat flow. Lithium was considered as the most
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
42
conservative trace element in this study, because the B concentration in several hot
spring waters were affected by phase separation.
The Li vs. Cl diagram shows that some of the fault-hosted hot springs have
similar high trends of Li enrichment as the volcano-hosted hot springs (Fig. 3.7). The
slopes of Li enrichment of the volcano-hosted hot springs are ~0.003, but ~0.003 and
~ 0.0001 for the fault-hosted hot springs. The similarity of high Li enrichment between
the fault-hosted and volcano-hosted geothermal fields could be an indication of the
same type of heat source, i.e, magmatic. Nevertheless, this assumption has to be
corroborated, because the trace element enrichment in hot springs can be generated
by several other processes, i.e., (1) high trace element concentration of the host rock,
(2) intermediate age and (3) high temperature of the geothermal system (Arehart et
al., 2003). The first point was ruled out, because the geothermal host rocks on Java
are basically identical, i.e., andesitic rocks. However, the second and the third points
were evaluated by considering the calculated reservoir temperatures. The similarity of
the high Li enrichment trend and the similarity high reservoir temperatures of fault-
hosted and volcano-hosted geothermal systems was considered an indication of a
magmatic heat source for both. Meanwhile, an intermediate age of a fault-hosted
geothermal system can be concluded when its reservoir temperature is lower than
that of a volcano-hosted geothermal system and its Li enrichment is as high that of a
volcano-hosted geothermal system. An intermediate age geothermal system has a
relatively higher trace element concentration due to the extended period of water-rock
interaction. Young geothermal systems have less time of water-rock interaction and
old geothermal systems have already leached most trace elments from their host
rocks, thus both systems have relatively low trace element concentrations (Arehart et
al., 2003). Considering those assumptions, the fault-hosted geothermal systems of
Cilayu, Batu Kapur, Maribaya and Cisolok should be heated by a magmatic heat
source. In contrast, such heat source was not likely for the Cikundul and Pakenjeng
fault-hosted geothermal systems. Hence, the high Li enrichment of these two fault-
hosted geothermal systems was caused by their intermediate age.
All of the low temperature and Li-poor fault-hosted geothermal systems,
Cikundul, Pakenjeng, Parangtritis and Pacitan, are located in the southern part of the
Java island. This area consists of the Tertiary volcanic belt, where volcanism ceased
in the last Paleogene. The volcanism then shifted northward forming the Neogene and
Quaternary volcanic belts in the central part of the island (Soeria-Atmadja et al.,
1994). Under those geological conditions, the heat source of those fault-hosted
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
43
geothermal system was likely similar to those described as ‘amagmatic’ heat source in
the Great Basin (USA) and Western Turkey (Faulds et al., 2010). An ‘amagmatic’ heat
source is a deep seated magma, which remained after volcanism ceased. The name
was used to distinguish this heat source from the shallow magmatic heat sources of
the Quaternary. As indicated by the exposure of the Cretaceous basement, the
southern part of the Java island underwent uplift and erosion due to the subduction of
Indo-Australia and Eurasian plates (Clements et al., 2009). As a result the crust
became thinner, which in turn increased the heat gradient, causing thermal circulation
of groundwater along faults, thus generating the fault-hosted geothermal systems.
The same heating mechanism was suggested for geothermal systems along the
Alpine fault, New Zealand (Allis and Shi, 1995; Shi et al., 1996).
Fig. 3.7. Li vs. Cl diagram of volcano-hosted and fault-hosted hot springs. Most of fault-hosted hot springs had similar trends of Li enrichment to those of volcano-hosted hot springs. (Symbols are as in Fig. 3.2).
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
44
III.4.4. Oxygen and hydrogen isotope considerations The deuterium and oxygen isotopic composition of hot springs can be used to
investigate their origin (Arnason, 1977; Craig, 1963; Craig et al., 1956; Giggenbach,
1978; Giggenbach et al., 1983; Majumdar et al., 2009; McCarthy et al., 2005; Pichler,
2005). All of the fault-hosted hot springs and most of the volcano-hosted hot springs
plotted close to the Local Meteoric Water Line (LMWL), indicating meteoric water as
the source of the hydrothermal fluids (Fig. 3.8).
Ten volcano-hosted geothermal waters, J9, J10, J19, J24, J26, J32, J51, J59
and J65, show stable isotope enrichment. This could be caused by either evaporation,
water-rock interaction, the input of magmatic fluids input or any combination of the
above (Craig, 1966; D'Amore and Bolognesi, 1994; Giggenbach and Stewart, 1982;
Ohba et al., 2000; Varekamp and Kreulen, 2000). Using the formulas from Gonfiantini
(1986) and Varekamp and Kreulen (2000), a theoretical evaporation line was
calculated for 90 °C lake temperature, 22 °C ambient temperature, 80% atmosphere
humidity and a δ18O of -6.9 ‰ and δ2H of -45 ‰ for meteoric water. Samples J10,
J26, J32 and J65 plotted close to the evaporation line although only J10, J32 and J65
were affected by evaporation, while the stable isotope enrichment in J26 was likely
caused by addition of andesitic water of Taran et al. (1989) and Giggenbach (1992).
The hot acid crater lakes of Kawah Kamojang (J9), Kawah Putih (J51) and
Kawah Sikidang (J59) have the heaviest isotope composition and plotted above the
field of andesitic water (Fig. 3.8). Connecting the hot crater lakes with their respective
meteoric water (δ18O = -6.9 ‰ and δ2H = -45 ‰) generates a line, which is flatter than
the evaporation line (Fig. 3.9). The slope of this line is close to the slope of other hot
acid crater lakes lakes, such as, Khusatsu-Shirane volcano (Ohba et al., 2000), Poas
volcano (Rowe Jr, 1994), Kelimutu (Varekamp and Kreulen, 2000) and Kawah Ijen
(Delmelle et al., 2000). Thus, the isotopic composition indicated that the hot crater
lake fluids likely underwent substantial evaporation and some reaction with magmatic
gas.
The presence of andesitic water in the geothermal systems was indicated for
samples J19 (Candi Songgoriti 2), J24 (Banyuasin) and J60 (Kawah Sileri), which plot
on or near the mixing line between local groundwater and andesitic water (Fig. 3.8).
However, Figure 3.10 only indicates andesitic water input for J19, J24 and J26, but
not for J60. The plot of J26 on the theoretical evaporation line in Figure 3.9 is caused
by its lighter stable isotope composition compared to the other hot springs with
andesitic water mixing. Andesitic water input was also found, for example, for the
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
45
Meager Creek (Clark et al., 1982), Larderello (D'Amore and Bolognesi, 1994), Geyser
(D'Amore and Bolognesi, 1994), Tongonan (Gerardo et al., 1993), El Chicon volcano
(Taran et al., 2008) and Tutum Bay (Pichler et al., 1999) geothermal systems (Fig.
3.9). In addition, the elevated Cl- concentration in samples J19, J24 and J26
corroborate the presence of andesitic water in those geothermal systems. In contrast
to these three samples, the presence of andesitic water could not be confirmed for
sample J60, due to its low Cl- concentration. Hence, similar to the J10, J32 and J65
hot springs, stable isotope enrichment of J60 should have been caused by
evaporation. The fact that J60 plots below the evaporation is likely due to a lighter
stable isotope composition of its meteoric source water compared to that of the
meteoric source water, which was used for the calculation of the theoretical
evaporation line.
Fig. 3.8. δ2H and δ18O compositions of cold and hot springs. LMWL and GMWL were taken from (Wandowo et al., 2001) and (Craig, 1961), respectively. All of fault-hosted and most of volcano-hosted hot springs had a meteoric water origin. Three stable isotope enrichments were identified, i.e., evaporation, combination of magmatic gases input with evaporation and andesitic water input. (Symbols are as in Fig. 3.2).
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
46
Fig. 3.9. δ2H and δ18O correlation lines between hot springs (gray filled) to their associated cold springs (open) from Java. The dashed lines are stable isotope correlation between thermal waters and their respective meteoric waters from other locations around the World.
III.5. Conclusions Based on the geological setting, two types of geothermal systems were
identified on Java, volcano-hosted and fault-hosted. Contribution of CO2 to the
geothermal fluid in volcano-hosted geothermal system, which was absent/minor in
fault-hosted geothermal system, led to a different water chemistry. Volcano-hosted hot
springs had higher HCO3- concentration and a higher Mg2+/Na+ ratio than fault-hosted
host springs. While the B concentration of fault-hosted hot springs was only affected
by water-rock interaction, volcano-hosted hot springs were influenced by phase
separation, water-rock interaction and groundwater mixing. Seawater addition was
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
47
identified in the Parangtritis, which was considered a fault-hosted geothermal system.
Calculated reservoir temperatures of volcano-hosted geothermal systems ranged from
125 to 338 °C and those of fault-hosted geothermal systems ranged from 74 to 299
°C.
Several geothermal systems on Java, although fault-hosted are likely heated by
shallow magmas, i.e., Batu Kapur, Maribaya, Cilayu and Cisolok. The addition of
volcanic fluids in Batu Kapur and Maribaya, which are located in the active
Quarternary volcanic belt, indicated that these two geothermal systems in fact are
volcano-hosted geothermal systems. Those fault-hosted geothermal fields that were
located in the old (Neogene) volcanic belt did not experience any addition of volcanic
fluids. Shallow magmas heat sources were not indicated in fault-hosted geothermal
systems of Cikundul, Pakenjeng, Pacitan and Parangtritis. Thus, deep seated magma
heat sources were suggested for those four geothermal systems.
Stable isotope enrichments were found in ten of the volcano-hosted geothermal
systems, but not in any of the fault-hosted geothermal systems. Stable isotope
enrichment due to evaporation was recognized in the Kawah Candradimuka and
Kawah Sileri, Kawah Hujan and Candi Gedong Songo geothermal systems. A
combination of intensive evaporation and magmatic gases input produced very heavy
stable isotopes in the hot acid crater lakes of the Kawah Kamojang, Kawah Sikidang
and Kawah Putih geothermal systems. The addition of substantial amounts of
Andesitic water to the geothermal fluid was observed in the Candi Songgoriti,
Banyuasin and Pablengan geothermal systems
Those finding reject the general assumption of a low energy potential of fault-
hosted geothermal systems, since although fault-hosted their heat source can be
magmatic as seen for several of the fault-hosted geothermal systems on Java. This
should give a new perspective for geothermal exploration on Java, where to date,
fault-hosted geothermal system were excluded from the geothermal energy
development program.
Acknowledgements A part of this study was supported by the Ministry of Energy and Mineral
Resources of Indonesia in a form of PhD scholarship grants. Thanks to Seto and
Maman for their help in the sampling campaign. TP acknowledges funding from the
German Research Foundation (DFG). M. Rowe and two anonymous reviewers are
thanked for their fruitful input and comments on the first version of this manuscript.
GEOTHERMAL SYSTEMS ON THE ISLAND OF JAVA, INDONESIA
48
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
49
IV. Boron isotope variations in geothermal systems on Java, Indonesia
Modified from: Budi Joko Purnomo a, Thomas Pichler a, & Chen-Feng You b
a Geochemistry & Hydrogeology, Department of Geosciences, University of
Bremen, Germany b Department of Earth Sciences, National Cheng Kung University, Tainan,
Taiwan, ROC
Submitted to:
Geochimica et Cosmochimica Acta
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
50
Abstract
This paper presents δ11B data for hot springs, hot acid crater lakes, geothermal brines
and a steam vent from Java, Indonesia with emphasis to investigate the difference
between boron (B) isotope compositions in volcano-hosted geothermal systems and
fault-hosted geothermal systems, as well as differences between sulfate crater lakes
and acid-chloride crater lakes. The possible seawater input into geothermal systems
and the mechanisms of B isotope fractionation were also investigated. The δ11B
values of hot springs ranged from -2.4 to +28.7 ‰ and hot acid crater lakes ranged
from +0.6 to +34.9 ‰. The δ11B and Cl/B values in waters from the Parangtritis and
Krakal geothermal systems indicated seawater input. The δ11B values of acid sulfate
crater lakes ranged from +5.5 to +34.9 ‰ and were higher than the δ11B of +0.6 ‰ of
the acid chloride crater lake. The heavier δ11B in the acid sulfate crater lakes was
caused by a combination of vapor phase addition and further enrichment due to
evaporation and B adsorption onto clay minerals. In contrast, the light δ11B of the acid
chloride crater lake was a result of continuous magmatic fluid input. The fluids in
volcano-hosted geothermal systems had a lighter δ11B signature than those in fault-
hosted geothermal systems. The faster ascent and the insignificant input of magmatic
fluids in the fault-hosted geothermal systems resulted in a light δ11B signature. In
contrast, the input of magmatic fluids and the slower ascent in volcano-hosted
geothermal systems were favorable for B isotope fractionation towards heavier
values.
Keywords:
Java, boron isotope, volcano- and fault-hosted geothermal systems, crater lake,
seawater input
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
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IV.1. Introduction Geothermal waters are known to have a large range of δ11B, from -9.3 to +44 ‰
(Aggarwal et al., 2000; Aggarwal et al., 1992; Barth, 1993; Leeman et al., 1990;
Musashi et al., 1988; Palmer and Sturchio, 1990; Vengosh et al., 1994b). This range
is caused by the δ11B signature of the geothermal host-rock, seawater input,
groundwater mixing and B isotope fractionation. Different sources of rocks were
identified producing a variety of δ11B in geothermal fluids at the Argentine Puna
Plateau (Kasemann et al., 2004). The heavy δ11B of seawater, +39.6 ‰ (Foster et al.,
2010), was successfully used to identify seawater components in geothermal waters
on Iceland (the Reykjanes and Svartsengi geothermal fields) and in Japan (the Izu-
Bonin arc, Kusatsu-Shirane area, and Kagoshima) (Aggarwal and Palmer, 1995;
Aggarwal et al., 2000; Kakihana et al., 1987; Millot et al., 2009; Musashi et al., 1988;
Nomura et al., 1982; Oi et al., 1993). Shallow groundwater dilution potentially
increases the δ11B composition of thermal waters (Palmer and Sturchio, 1990; Yuan
et al., 2014). Fractionation of B isotopes in thermal waters occurred due to
adsorption/incorporation of B onto clay minerals and iron oxide (Lemarchand et al.,
2007; Palmer et al., 1987; Schwarcz et al., 1969; Spivack and Edmond, 1987;
Vengosh et al., 1991b), calcite (Hemming and Hanson, 1992; Vengosh et al., 1991a)
and evaporite minerals (Agyei and McMullen, 1968; McMullen et al., 1961; Oi et al.,
1989; Swihart et al., 1986; Vengosh et al., 1992). These processes enrich the 10B
isotope in the solid phases and thus increase the δ11B of thermal waters. Thermal
waters which condensated from the vapor phase of a geothermal system are
potentially enriched in δ11B as a consequence of 11B fractionation into the vapor
phase. The δ11B enrichment during this process was generally considered
insignificant (Kanzaki et al., 1979; Leeman et al., 1992; Nomura et al., 1982; Spivack
et al., 1990; Yuan et al., 2014), but potentially should not be neglected in vapor-
dominated geothermal systems.
Java has many hot springs, steam vents, mud pools, hot acid crater lakes and
altered grounds. Purnomo and Pichler (2014) divided the geothermal systems into
fault-hosted and volcano-hosted geothermal systems, based on their location, either
in a volcanic complex or a fault zone. While in fault-hosted geothermal systems deep
circulating groundwater is simply heated, volcano-hosted geothermal systems seem
to be more complex due to the addition of magmatic fluids to the geothermal water
(Alam et al., 2010; Purnomo and Pichler, 2014).
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
52
The focus of previous B isotope studies in geothermal systems were generally
the mechanisms of B isotopes fractionation, while in this paper B isotope signatures
were rather used as tracers to investigate contrasting systems, i.e., (1) volcano-
hosted vs. fault-hosted geothermal systems, (2) acid sulfate vs. acid chloride crater
lakes and (3) meteoric vs. seawater influenced.
IV.2. Sampling locations Thermal water samples for this study were taken from 21 geothermal systems
on Java consist of 15 volcano-hosted and 6 fault-hosted geothermal systems (Fig.
4.1). All fault-hosted geothermal systems are distributed in the Tertiary volcanic belt.
The Cikundul and the Parangtritis geothermal systems are hosted in the major fault
zones of Cimandiri and Opak, respectively (Effendi et al., 1998; Rahardjo et al., 1995).
Other fault-hosted geothermal systems, i.e., Cisolok, Cilayu, Pakenjeng and Krakal,
are hosted in minor fault zones (Alzwar et al., 1992; Asikin et al., 1992; Silitonga,
1973; Sujatmiko and Santosa, 1992). Meanwhile, all volcano-hosted geothermal
systems on Java are within the Quaternary volcanic complex. The Kamojang, Darajat
and Wayang-Windu are located in the Kendang volcanic complex (Rejeki et al., 2005).
The Patuha geothermal system is located in the flat volcanic highland of the Patuha
volcano (Layman and Soemarinda, 2003). The Sari Ater geothermal system is hosted
by the Tangkuban Prahu volcano, Cileungsing by the Tampomas volcano and
Segaran by the Lamongan volcano. The Gucci and Baturaden geothermal systems
are hosted by the Slamet volcano. Another single volcano, the Lawu volcano is
hosting the Lawu geothermal system. The Songgoriti and Padusan geothermal
systems are located in the Arjuna-Weilrang volcano complex and the Dieng
geothermal system is located in the Dieng caldera.
IV.3. Results The temperature, pH, Cl-, B, HCO3
-, Fe data and δ11B values are reported in
Table 4.1. Except for δ11B, this data is from Purnomo and Pichler (2014). The
physicochemical data and δ11B values of new samples from the Kawah Kreta steam
vent (J73), the Krakal hot spring (J74), the Kawah Domas acid sulfate crater lake
(J72) and two geothermal brines of well AFT-28 (J70) and PAD-7C (J71) from the
Dieng geothermal field are reported separately in Table 4.2.
Fig.
4.1
. The
sam
plin
g lo
catio
n of
geo
ther
mal
fiel
ds o
n Ja
va, i
.e.,
(1)
Cis
olok
, (2)
Cik
undu
l, (3
) Ba
tu K
apur
, (4)
Tan
gkub
an P
arah
u, (
5)
Tam
pom
as, (
6) P
atuh
a, (7
) Pan
gale
ngan
, (8)
Dar
ajat
, (9)
Kam
ojan
g, (1
0) C
ipan
as, (
11) C
iaw
i, (1
2) C
ilayu
, (13
) Pak
enje
ng, (
14) S
lam
et,
(15)
Kra
kal,
(16)
Die
ng, (
17)
Kalia
nget
, (18
) Pa
rang
tritis
, (19
) La
wu,
(20
) A
rjuna
-Wel
irang
and
(21
) S
egar
an. G
eolo
gica
l stru
ctur
es a
nd
volc
anic
bel
ts w
ere
base
d on
Ham
ilton
(197
9), S
iman
djun
tak
and
Bar
ber (
1996
), H
offm
ann-
Rot
he e
t al.
(200
1) a
nd S
oeria
-Atm
adja
et a
l. (1
994)
. Mod
ified
from
Pur
nom
o an
d P
ichl
er (2
014)
.
) t, d l.
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
54
Table 4.1. Temperature, pH, Cl-, HCO3-, B and Fe data and δ11B of hot springs, hot
acid crater lakes and cold springs from Java. Sample Location Geo. Temp. pH Cl HCO3 B Fe δ11B
ID Type °C mg/L ‰
Hot springs J2 Baturaden (Slamet volc.) V 52 6.9 777.3 722.2 4.3 3.0 5.9
J4 Ciawi V 53 6.7 165.3 976.0 6.8 0.6 -0.7
J7 Cipanas V 48 6.3 119.0 383.1 2.4 0.01 4.6
J10 Kawah Hujan (Kamojang) V 95 4.9 7.2 22.0 4.6 0.1 2.3
J13 Padusan (Arjuna-Welirang volc.) V 48 6.5 246.2 1104.1 5.0 1.7 12
J18 Songgoriti (Arjuna-Welirang volc.) V 46 6.3 1303.5 1378.6 50.6 6.2 7.1
J21 Segaran (Lamongan volc.) V 45 6.5 550.5 1625.0 21.4 0.03 6.7
J24 Banyuasin (Lawu volc.) V 38 6.1 5948.7 835.7 93.2 9.4 8.1
J28 Kondo (Lawu volc.) V 36 6.3 4382.4 1634.8 55.6 1.9 4.5
J30 Ngunut (Lawu volc.) V 33 6.4 904.9 878.4 17.4 0.02 1.3
J34 Gucci (Slamet volc.) V 53 7.5 52.6 556.3 7.2 0.2 7.8
J36 Sari Ater (Tangkuban Parahu volc.) V 47 2.0 822.5 <dl 2.1 21.6 2.4
J38 Batu Kapur V 41 6.4 312.7 1085.8 3.1 1.3 -2.4
J39 Cibolang (Pangalengan) V 69 7.1 24.0 219.6 6.5 0.01 0.4
J42 Pakenjeng F 60 7.4 126.0 40.3 7.2 0.06 0.0
J44 Cilayu F 70 8.1 1387.2 372.1 58.2 0.1 0.2
J47 Kalianget V 40 6.5 424.8 732.0 3.7 1.1 8.9
J48 Cikundul F 51 7.8 180.2 61.0 10.7 0.01 9.3
J49 Cisolok F 102 8.1 305.6 129.3 3.6 0.01 -0.7
J54 Cileungsing (Tampomas) V 49 7.1 757.2 732.0 5.3 1.8 -1.0
J55 Darajat V 60 2.8 13.3 < dl 7.0 8.6 12.8
J58 Parangtritis F 39 7.6 6184.5 43.9 9.5 0.2 24.8
J60 Kawah Sileri (Dieng) V 57 6.3 77.7 266.0 6.7 0.02 0.3
J61 Pulosari (Dieng) V 54 6.2 330.6 183.0 6.4 0.1 4.6
J64 Bitingan (Dieng) V 60 7.3 21.5 329.4 2.1 0.01 3.2
Hot acid crater lakes
J9 Kawah Kamojang (Kamojang) V 40 2.9 12.8 <dl 1.4 4.9 9.3
J51 Kawah Putih (Patuha) V 33 1.0 8084.2 <dl 94.4 35.3 0.6
J59 Kawah Sikidang (Dieng) V 87 2.5 14.4 <dl 73.3 7.6 34.9
Cold springs J5 Ciawi 25 6.5 9.1 97.6 0.01 0.01 nm
J15 Cangar (Arjuna-Welirang volc.) 23 6.3 22.0 115.9 0.3 0.01 11.1
J67 Bimo Lukar (Dieng) 18 6.2 17.3 7.3 0.01 0.01 nm
*V= volcano-hosted, F= fault-hosted, nm= not measured, <dl= below detection limit
Tabl
e 4.
2. P
hysi
coch
emic
al d
ata
and δ11
B va
lues
of t
he K
raka
l hot
spr
ing,
the
Kaw
ah D
omas
cra
ter l
ake,
the
Kaw
ah K
reta
ste
am v
ent
and
geot
herm
al b
rines
of t
he D
ieng
geo
ther
mal
fiel
d.
Sam
ple
Loca
tion
Geo.
Te
mp.
pH
O
RP
Cond
TD
S Ca
M
g N
a K
Cl
HCO
3 SO
4 Si
Sr
Al
B
Mn
Fe
δ11B
ID
Type
°C
m
V uS
/cm
m
g/L
‰
Hot s
prin
gs
J74
Krak
al, K
ebum
en
F 39
8.
2 -8
6 21
470
2017
0 20
82.9
<d
l 24
36.4
3.
8 86
71.8
31
.2
21.6
<d
l 8.
8 0.
4 14
.5
< dl
< dl
28
.7
Hot a
cid
crat
er la
kes
J72
Kaw
ah D
omas
V
85
1.8
94
6935
56
08
38.1
8.
5 17
.9
10.9
<d
l <d
l 20
40
86.4
0.
1 70
.9
2.7
1.2
17.4
5.
5
(T.
Para
hu)
Geo
ther
mal
brin
es
J70
GW A
FT-2
8 (D
ieng
) V
150
6.3
-196
59
500
6400
0 89
3.5
<dl
1053
6.3
3016
.5
2224
2.5
39.0
<d
l 39
4.3
8.9
<dl
593.
6 12
.21.
3 0.
3
J71
GW P
AD-7
C (D
ieng
) V
150
5.8
-117
37
110
3850
0 37
1.6
<dl
6504
.5
1933
.4
1338
4.0
14.6
38
.8
457.
9 5.
6 <d
l 26
2.5
8.2
2.3
0.0
Stea
m v
ent
J73
Kaw
ah K
reta
(K
amoj
ang)
V
nm
nm
nm
nm
nm
2.0
<dl
0.8
<dl
<dl
nm
4.7
<dl
<dl
0.14
3.
5 <d
l<d
l 3.
8
*V=
volc
ano-
host
ed, F
= fa
ult-
host
ed, n
m=
not m
easu
red,
<dl
= be
low
det
ectio
n lim
it
11B
‰ 8.7
5.5
0.3
0.0
3.8
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
56
The B concentrations of hot springs and acid crater lakes had a similar large
range, from 2.1 to 93.2 mg/L and from 1.4 to 94.4 mg/L, respectively. The two
geothermal brines from the Dieng geothermal field had B contents of 262.5 and 593.6
mg/L. The Kawah Kreta steam vent (J73) had a B content of 3.5 mg/L, which is
enriched relative to its Cl- concentration, which was below detection. The B
concentration in the acid sulfate crater lakes ranged from 1.4 to 73.3 mg/L, while the
Kawah Putih (J51) acid chloride crater lake had a slightly higher value of 94.4 mg/L.
The δ11B of the Cangar cold spring (J15) was +11.1 ‰, within the range of the
thermal waters of -2.4 to +34.9 ‰, but heavier than the Kawah Kreta steam vent of
+3.8 ‰ and the Dieng geothermal brines of 0 to 0.3 ‰. In accordance with the B
concentration, the hot springs and hot crater lakes had relatively similar ranges of δ
11B, i.e., -2.4 to +28.7 ‰ and +0.6 to 34.9 ‰, respectively. In contrast to the B
content, acid sulfate crater lakes had relatively heavy δ11B values from +5.6 to 34.9
‰compared to the acid chloride crater lake, which had a value of +0.6 ‰. The fault-
hosted and volcano-hosted geothermal systems had a relatively similar large ranges
of δ11B, i.e., -2.4 to +28.7 ‰ and -1.0 to +34.9 ‰, respectively. The J74 and J58 fault-
hosted hot springs as well as J59 acid sulfate crater lake had a δ11B value close to
that of seawater, i.e., +39.6 ‰ (Foster et al., 2010). The δ11B of the thermal waters
were poorly correlated with their temperature, pH, Cl-, B, HCO3- and Fe (Fig. 4.2).
Thermal waters with temperatures above 65 °C generally had relatively light δ11B
values, close to 0 ‰, while below this temperatures most of the thermal waters were
δ11B enriched (Fig. 4.2a). This indicates more significant of B isotope enrichment at
low temperatures, as suggested by others, e.g., Palmer et al. (1987) and Aggarwal
and Palmer (1995).
IV.4. Discussion IV.4.1. Boron in thermal waters and seawater input
Following the procedure of Arnorsson and Andresdottir (1995) the B content of
thermal waters on Java was used to identify steam separation for J9, J10, J34, J39,
J55, J59 and J64; andesitic host-rock leaching for most of the thermal waters;
seawater input for J58; and B adsorption for J2, J36, J38, J47, J49 and J54 (Purnomo
and Pichler, 2014). The Krakal (J74) fault-hosted hot spring plotted close to the
seawater line, similar to J58, indicating seawater input, while the Kawah Domas acid
sulfate crater lake (J72) plotted in the vapor phase separation and the geothermal
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
57
brines from Dieng, J70 and J71, plotted close to the andesitic rock leaching line (Fig.
4.3).
Fig. 4.2. a) δ11B vs. T, b) δ11B vs. pH, c) δ11B vs. Cl, d) δ11B vs. B, e) δ11B vs. HCO3 and f) δ11B vs. Fe diagrams of thermal waters on Java show poor correlations of δ11B with T, pH, Cl-, B, HCO3
- and Fe. The δ11B vs. T indicates generally lower δ11B enrichment at temperatures above 65 °C.
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
58
Fig. 4.3. Cl vs. B diagram of thermal waters (modified from Purnomo and Pichler, 2014). Thermal waters underwent seawater input, B depletion, steam phase separation and andesitic rock leaching. Seawater input is indicated for two fault-hosted geothermal systems, J58 and J74, while most thermal waters were resulted by andesitic rock leaching.
The boron isotope composition can be a powerful tool to detect seawater input
into geothermal systems, as demonstrated for two geothermal fields on Iceland (the
Reykjanes and Svartsengi) and three areas in Japan (the Izu-Bonin arc, Kusatsu-
Shirane area, and Kagoshima) (Aggarwal and Palmer, 1995; Aggarwal et al., 2000;
Kakihana et al., 1987; Millot et al., 2009; Musashi et al., 1988; Nomura et al., 1982; Oi
et al., 1993). A similar approach was used to investigate seawater input in the fault-
hosted thermal waters at Krakal (J74) and Parangtritis (J58). These two thermal
waters plot close to the mixing line between seawater and groundwater (Fig. 4.4),
which proves the presence of seawater. The acid sulfate crater lake of Kawah
Sikidang (J59) also had a heavy δ11B of +34.9 ‰. However, this acid sulfate crater
lake had a Cl/B ratio five magnitudes lower than seawater (Fig. 4.4), thus seawater
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
59
input could be excluded. The fluid source of Kawah Sikidang, which is the deep
reservoir of the Dieng geothermal field, J70 and J71, also showed the absence of
seawater input. The geothermal brines had a light δ11B of approximately 0 ‰ and Cl/B
ratios two magnitudes lower than seawater (Fig. 4.4). If there was seawater input
during the vapor phase ascent, the resulting thermal water should have a higher Cl-
content corresponding to the vapor/seawater mixing ratio.
Fig. 4.4. δ11B vs. B/Cl diagram of thermal waters, groundwater and geothermal brines. Seawater input is confirmed for two fault-hosted thermal waters, J58 and J74, but not for the Kawah Sikidang (J59) acid sulfate crater lake. The δ11B value of +39.6 ‰ for seawater (Foster et al., 2010) and B/Cl ratio of the Indian Ocean (Purnomo and Pichler, 2014) are used in this diagram.
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
60
IV.4.2. The δ11B of acid sulfate and acid chloride crater lakes Two species of B, B(OH)3 and B(OH)4
- exist in natural waters and only the
former species is present at a pH of less than 7 (Dickson, 1990; Xiao et al., 2013).
B(OH)3 species produce a higher degree of B isotope fractionation than B(OH)4-,
hence B isotope fractionation is stronger at low pH (Palmer et al., 1987). The four acid
crater lakes, Kawah Kamojang (J9), Kawah Domas (J72), Kawah Putih (J51) and
Kawah Sikidang (J59) had low pH values ranging from 1 to 2.9 and thus were
favorable for B isotope fractionation.
The B/Cl ratios indicate that, except for Kawah Putih, the three other crater
lakes were formed due to condensation of the geothermal vapor phase (Fig. 4.3). The
δ11B of vapor phase is generally enriched by up to 4 ‰ relative to the remaining liquid
phase (Kanzaki et al., 1979; Leeman et al., 1992; Nomura et al., 1982; Spivack et al.,
1990; Yuan et al., 2014). The difference in δ11B between the Dieng geothermal brines
and the Kawah Kreta vapor phase was approximately 3.8 ‰ confirming condensation
as the main process of fractionation. Assuming steam condensation from a similar
vapor phase as Kawah Kreta, the Kawah Domas, Kawah Kamojang and Kawah
Sikidang crater lakes were enriched in δ11B by 1.7 ‰, 5.5 ‰ and 31.1 ‰,
respectively. The Kawah Kreta steam vent had a B concentration of 3.5 mg/L, which
was slightly higher than Kawah Kamojang with 1.4 mg/L and Kawah Domas with 2.7
mg/L. Smith et al. (1987) reported for the Geyser geothermal field (USA) that the
water of the vapor trap was approximately 50 times enriched in B concentration
compared to the steam phase. Accordingly, an acid crater lake originated from vapor
phase condensation is expected to have a higher B concentration than the vapor
phase. Therefore, the low B concentration of the Kawah Kamojang and Kawah
Domas crater lakes in comparison to the vapor phase should indicate additional
processes which lowered the B concentration. This could have been caused either by
precipitation of a B-rich mineral phase or adsorption by clay minerals. Both processes
reduce the B concentration of the remaining water and thus enrich 11B due to
adsorption of 10B by the solid phases (Agyei and McMullen, 1968; McMullen et al.,
1961; Oi et al., 1989; Palmer et al., 1987; Schwarcz et al., 1969; Spivack and
Edmond, 1987; Swihart et al., 1986; Vengosh et al., 1991b; Vengosh et al., 1992).
Palmer and Sturchio (1990) reported more significant B isotope fractionation at low
temperature and thus the high temperature of Kawah Domas (85 °C) should allow
only little B isotope fractionation, causing the lighter δ11B values compared to Kawah
Kamojang, where the temperature was 40 °C. Meanwhile, the heavy δ11B value in
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
61
Kawah Sikidang, which was by 31 ‰ heavier than the vapor phase, cannot be
explained exclusively by steam phase separation. Other mechanisms must have
produced such a B-rich and heavy δ11B signature in this acid sulfate crater lake.
Apart from the heavy δ11B, the Kawah Sikidang crater lake had a noticeable B
enrichment of 73.3 mg/L. This value is higher than that of the other acid sulfate crater
lakes and only comparable to the acid chloride crater lake Kawah Putih, which had a
B concentration of 94.4 mg/L. However, Kawah Putih was Cl-rich (8084 mg/L) and its
δ11B was +0.6 ‰, whereas Kawah Sikidang in contrast was Cl-poor (14.4 mg/L) with a
δ11B of +34.9 ‰. Delmelle and Bernard (1994) explained that the chemical
composition of an acid chloride crater lake is the result of condensation and oxidation
of magmatic gases, such as, SO2, H2S and HCl, upon contact with oxygenated
groundwater followed by water-rock interaction. Fumaroles in Japan (Kanzaki et al.,
1979; Nomura et al., 1982) and Vulcano Island, Italy (Leeman et al., 2005), for
example, had light δ11B, representing the magmatic fluids. Accordingly, the vapor
phase of Kawah Putih should have been derived from magmatic fluids and thus
inherited the light δ11B of an andesitic island arc magma of -2.3 to +3.5 ‰ (Palmer,
1991). After discharge at the surface, B isotope enrichment was moreless absent
though the water was very acidic (pH = ~1) and relatively cold (T = 32.9 °C), probably
was caused by continuous magmatic fluid input. The crater lake had a dimension of
approximately 350 x 290 m2 (Sriwana et al., 2000) and the ambient air temperature
was 16 °C at the time of sampling, that the fluid was cooled down rapidly.
The main mechanism that produces heavy δ11B (+34.9 ‰) of Kawah Sikidang
probably was evaporation. Vengosh et al. (1992) reported that during the latest stage
of evaporation the δ11B of seawater could be enriched up to 30 ‰ due to incorporation
of 10B into evaporate minerals, e.g., halite. Kawah Sikidang had a temperature of 87
°C without outflow to the nearby river, hence excessive evaporation could be
expected. The condition is contrary to the Kawah Domas acid sulfate crater lake
which had a temperature of 85 °C but the water drained to the adjacent river and thus
the effect of evaporation was less significant than at Kawah Sikidang. Evaporation
should have lowered the B concentration of Kawah Sikidang due to incorporation of B
into evaporites. Therefore, the elevated B content of 73.3 mg/L of the crater lake has
to be sustained by constant addition of subsurface B. The reservoir vapor phase
probably had an initial B content comparable to the Kawah Kreta steam vent of 3.5
mg/L, however, in the shallow depth the vapor was B enriched due to interaction with
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
62
B-rich minerals. This assumption was supported by the elevated B concentration of
altered ground surrounding Kawah Sikidang which was up to 25.5 mg/Kg, compared
to the less than 0.02 mg/Kg of unaltered ground in Dieng.
IV.4.3. Processes affecting the δ11B value of thermal waters
Mixing with groundwater in the shallow subsurface prior to discharge leads to a
heavier δ11B signature, because groundwater has generally a heavier δ11B signature
than thermal water (Palmer and Sturchio, 1990; Vengosh et al., 1994b). A binary
mixing model can be applied to estimate the effect of groundwater mixing for the B
concentration and δ11B signature of thermal waters (Vengosh et al., 1991b; Yuan et
al., 2014). To set up the model, the groundwater of Cangar (J15) and the geothermal
brine from the Dieng geothermal field (J70) were used as end members. Groundwater
dilution was evident for eight thermal waters, i.e., J10 (Kawah Hujan), J30(Ngunut),
J36 (Sari Ater), J39 (Cibolang), J42 (Pakenjeng), J44 (Cilayu), J60 (Kawah Sileri) and
J64 (Bitingan) (Fig. 4.5). However, most of the thermal waters plot above the mixing
line and four thermal waters, J4 (Ciawi), J38 (Batu Kapur), J49 (Cisolok) and J54
(Cileungsing), plot below the mixing line (Fig. 4.5), indicating that the B concentration
and δ11B signature in those waters is more complex than just groundwater mixing.
Those samples plotting above ther mixing line are relatively enriched, which could
have been caused by adsorption, phase separation, water-sediment interaction or any
combination of these processes (Hemming and Hanson, 1992; Lemarchand et al.,
2007; Palmer et al., 1987; Schwarcz et al., 1969; Spivack and Edmond, 1987;
Vengosh et al., 1991a; Vengosh et al., 1991b; Vengosh et al., 1992; Xiao et al., 2013;
Yuan et al., 2014). The four thermal waters plotting below the mixing line either had
initially a light δ11B signature or desorbed B from solid phases. Although J36, J10, J39
and J64 plotted close to the groundwater mixing line (Fig. 4.5), groundwater mixing
cannot be the only process responsible for their final δ11B composition. Based on their
B/Cl ratios, J36 was relatively B depleted, while J10, J39 and J64 were relatively B
enriched (Fig. 4.3). The B concentration in J36 was lowered due to adsorption onto
clay minerals, which at the same time increased the δ11B value to +2.4 ‰. Since J36
was of the acid chloride type with a pH of 2, it likely had an initial δ11B value similar to
the geothermal brines of approximately 0 ‰. In contrast, the higher B/Cl ratios of J10,
J39 and J64 compared to the andesitic rock were a result of steam phase separation
(Fig. 4.3). Thermal waters generated by condensation of a vapor phase in
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
63
Fig. 4.5. δ11B vs. B diagram of thermal waters and geothermal brines with a theoretical binary mixing line of groundwater (J15) and geothermal brine (J70). Groundwater dilution is dominant for J10, J30, J36, J39, J42, J44, J60 and J64 thermal waters.
a geothermal system generally undergo δ11B enrichment of up to 4 ‰ (Leeman et al.,
1992; Spivack et al., 1990). J10 and J64 had a δ11B of +2.3 ‰ and +3.2 ‰,
respectively, which were close to the Kawah Kreta steam vent of +3.8 ‰. The slightly
lower δ11B of those two thermal waters compared to the steam vent could be due to
addition of 10B from carbonate mineral dissolution, because J10 and J64 were under
saturated with respect to carbonate minerals (Table 4.3). In the case of J64
desorption of B from iron oxide could have been possible as well, since iron oxides
were also undersaturated (Table 4.3). In comparison, J39 was saturated with respect
to goethite, carbonate and clay minerals and its temperature of 68.9 °C would
dampen B isotope fractionation after discharge. Therefore, the light δ11B of J39 must
be derived from its vapor phase leading to the conclusion that B isotope fractionation
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
64
Table 4.3. Saturation indices with respect to carbonate, clay and goethite minerals were calculated using PHREEQC (Parkhurst and Appelo, 1999).
Sample Location Saturation Index
ID Calcite Dolomite Strontianite Illite Smectite Goethite
J2 Baturaden (Slamet volc.) 0.64 2.74 -0.20 - - 1.56
J4 Ciawi 0.33 2.12 -0.61 - - 1.81
J7 Cipanas -0.81 -0.02 - - - 2.04
J10 Kawah Hujan (Kamojang) -4.27 -7.78 -5.49 2.12 -2.54 1.40
J13 Padusan (Arjuna-Welirang) 0.22 1.81 -0.71 - - 1.50
J18 Songgoriti (Arjuna-Welirang) 0.03 1.45 -0.31 6.96 3.48 0.62
J21 Segaran (Lamongan volc.) 0.15 2.13 -0.33 - - -1.31
J24 Banyuasin (Lawu volc.) -0.28 0.38 -0.08 - - 1.16
J28 Kondo (Lawu volc.) -0.25 0.51 0.37 - - -1.03
J30 Ngunut (Lawu volc.) -0.27 0.33 -0.08 - - 0.77
J34 Gucci (Slamet volc.) 0.73 3.06 -0.21 4.03 4.48 4.79
J74 Krakal, Kebumen 1.14 - 0.39 - - -
J38 Batu Kapur -0.14 1.19 -0.26 - - 0.29
J39 Cibolang (Pangalengan) 0.35 1.89 -0.69 1.40 2.29 4.74
J42 Pakenjeng 0.09 - -0.65 - - 3.82
J44 Cilayu 1.32 3.42 0.92 -3.13 1.81 7.01
J47 Kalianget -0.12 1.31 -1.15 - - 0.81
J48 Cikundul -0.14 -0.36 - 3.69 2.02 -0.01
J49 Cisolok 1.18 2.62 0.42 - - 4.49
J54 Cileungsing (Tampomas) 0.54 2.39 0.51 - - 2.01
J58 Parangtritis 0.76 0.64 0.01 - - 4.00
J60 Kawah Sileri (Dieng) -0.80 -0.60 -1.75 3.02 -0.14 -4.36
J61 Pulosari (Dieng) 0.08 1.20 -0.64 0.50 0.80 3.78
J64 Bitingan (Dieng) -1.28 -1.61 -2.07 4.03 -0.98 -4.27
during vapor phase separation must have been insignificant. The isotope fractionation
of B is considered unimportant at reservoir temperatures above 400 °C (Leeman et
al., 1992; Spivack et al., 1990) or high reservoir pressure (Liebscher et al., 2005). The
reservoir temperature of J39 ranged from 221 to 323 °C, similar to J10 and J64 of
approximately ranging from 220 to 350 °C (Purnomo and Pichler, 2014; Sudarman et
al., 1995), hence temperature cannot be responsible for the light δ11B of the J39 vapor
phase. Meanwhile, the reservoir pressure of J39 (Wayang-Windu geothermal field)
was recorded up to 84 bar (Mulyadi and Ashat, 2011), higher than J10 (Kamojang
geothermal field) and J64 (Dieng geothermal field) of up to 36 and 40 bar, respectively
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
65
(Bachrun et al., 1995; Zuhro, 2004). Therefore, the light δ11B of the J39 vapor phase
was more likely caused by high reservoir pressure.
Twelve thermal waters, J2, J7, J13, J18, J21, J24, J28, J34, J47, J48, J55 and
J61, were δ11B enriched, thus plotted above the mixing line (Fig. 4.5). The increasing
δ11B in a hot spring where adsorption/coprecipitation of B occurs should coincide with
a decrease of the B/Cl ratio. However, as seen in the Cl vs. B diagram only J2
underwent B depletion, while J55 plotted in the steam separation group and J7, J13,
J18, J21, J24, J28, J34, J48 as well as J61 plotted close to andesitic rock leaching
(Fig. 4.3). Compared to the Kawah Kreta steam vent, J55 was δ11B enriched by 9 ‰,
a result of B adsorption onto clay minerals due to its acidic condition (pH = 2.8). This
process lowered the B/Cl ratio of J55, however, the thermal water still plotted in the
steam separation group due to its initial higher B/Cl ratio, formed by the high
temperature geothermal system. Meanwhile, J7, J13, J18, J21, J24, J28, J34, J48
and J61 plot close to the andesitic line, which might be caused by Cl- depletion during
ascent or initially higher B/Cl ratios. In a mud volcano, for example, Cl- depletion was
considered due to NaCl filtration by clay minerals during clay dehydration (Dahlmann
and Lange, 2003; Dia et al., 1999; Hensen et al., 2004; Kastner et al., 1991; You et
al., 2004). This mechanism, however, is unlikely in a geothermal system, clay
minerals along the fluid pathway should be hydrated instead. Vapor phase separation
and reaction with B-rich minerals potentially increase the B/Cl ratio of a thermal water.
However, the former could be ruled out because those nine thermal waters had higher
Cl- contents compared to the thermal waters generated by steam separation.
Therefore, it could be suggested that they underwent B enrichment due to interaction
with B-rich minerals in the subsurface. The minerals supplied B into the ascending
thermal water and thus increased the B/Cl ratio. Post discharge at the surface, the
water underwent B adsorption/coprecipitation that reduced the B/Cl ratio but
increased the δ11B.
In order to identify the mineral phases that potentially adsorb/desorp B in hot
springs, the saturation indices of carbonate, clay and iron oxide minerals were
calculated using PHREEQC of Parkhurst and Appelo (1999) (Table 4.3). Lemarchand
et al. (2007) reported a higher B isotope fractionation during adsorption onto iron
oxide than onto clay and carbonate minerals. The saturation index calculation
indicates that B coprecipitation with goethite could be expected for most of the hot
springs, but not for J21, J28, J74, J48, J60 and J64 (Table 4.3). In hot springs J21
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
66
and J74 B likely coprecipitated with calcite, in J48, J60 and J64 B adsorbed onto clay
minerals and in J28 it coprecipitated with dolomite and strontianite.
IV.4.4. The δ11B of fault-hosted and volcano-hosted geothermal systems
The magmatic fluid input in volcano-hosted geothermal systems, which is
absent/minor in fault-hosted geothermal systems, distinguishes their chemical
characteristics (Purnomo and Pichler, 2014). The presences of magmatic gases (CO2,
H2S, HCl, SO2) in volcano-hosted geothermal systems produces reactive thermal
waters, favorable for clay, carbonate and sulfate mineral precipitation. These minerals
then can preferentially adsorb 10B over 11B causing the δ11B of thermal waters to
increase. In addition, the longer flow path and thus ascent time of volcano-hosted
thermal waters compared to fault-hosted geothermal systems enhances B isotope
fractionation. Considering these two factors, volcano-hosted thermal waters should
have a heavier δ11B signature than fault-hosted thermal waters. By excluding thermal
waters influenced by seawater and acid crate lakes, this assumption could be
confirmed.
With the exception of J48 (Cikundul), the fault-hosted thermal waters had δ11B
values ranging from -0.7 to +0.2 ‰, while volcano-hosted thermal waters were
generally heavier and ranged from -2.4 to +12.8 ‰. Figure 4.6 indicates that most of
the volcano-hosted thermal waters were δ11B enriched, while for the fault-hosted such
indication was only found in J48. Five volcano-hosted thermal waters, J4, J38, J39, 54
and J60, which had light δ11B signatures were saturated with carbonate minerals and
goethite (Table 4.3). The addition of 10B to those thermal water due to mineral
dissolution can be dismissed and thus sources of those thermal waters should have
had a lighter δ11B signature than the geothermal reservoir at Dieng. Meanwhile, the
relatively heavy δ11B (+9.3 ‰) and high B value (10.7 mg/L) of the fault-hosted hot
spring J48 are considered a result of water-sediment interaction, since this hot spring
was located at the bottom of the Cimandiri river. The relatively high B value and a
temperature of 51 °C indicate insignificant mixing with river water prior to sampling.
IV.5. Conclusions Hot springs and hot crater lakes on Java had a range of δ11B from -2.4 to +34.9
‰, relatively similar to other geothermal systems in the World (Barth, 1993). Seawater
input was detected for two fault-hosted geothermal systems, Parangtritis and Krakal,
and is considered a source of the heavy δ11B. The Kawah Sikidang acid sulfate crater
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
67
Fig. 4.6. δ11B vs. B diagram indicates most volcanic-hosted thermal waters were δ11B enriched, while δ11B enrichment for fault-hosted thermal waters were moreless absent.
lake had a δ11B indicating seawater addition to its geothermal source, although the
low Cl/B ratio excluded seawater as a source for 11B.
Two types of hot acid crater lakes, Cl-poor (Cl- < 15 mg/L) and Cl-rich (Cl- =
8084 mg/L), showed a contrasting δ11B composition. The Cl-poor crater lakes had
relatively heavy δ11B values ranging from +5.5 to 34.9 ‰, while the Cl-rich lake had a
lighter δ11B of +0.6 ‰, similar to the Dieng geothermal brines. The heavier δ11B
values of the acid sulfate crater lakes are a combination of vapor phase separation,
evaporation after discharge and preferential adsorption of 10B by clay minerals. The
heavy δ11B (+34.9 ‰) of the Kawah Sikidang acid sulfate crater lake was mainly
caused by evaporation, while its B content of 73.3 mg/L was due to water-rock
interaction with B-rich minerals in the shallow subsurface. Meanwhile, the light δ11B of
B O R O N I S O T O P E V A R I A T I O N S I N G E O T H E R M A L S Y S T E M S O N J A V A , I N D O N E S I A
68
the Kawah Putih acid chloride crater lake was a result of continuous magmatic gases
input.
The initial division of the geothermal systems on Java into fault-hosted and
volcano-hosted (Purnomo and Pichler, 2014) was confirmed by their different δ11B
signatures. In general, fault-hosted systems had lighter δ11B values ranging from -0.7
to +0.2 ‰, while the volcano-hosted systems had heavier δ11B values of above 1 ‰.
The comparably faster ascent to the surface and the absence of magmatic fluids in
fault-hosted geothermal systems should not facilitate B isotope fractionation, hence
the lower δ11B values. In volcano-hosted geothermal systems, the ascent time is
relatively longer and the condensation of magmatic gas, as well as the formation of
clay, carbonate, sulfate and iron oxide minerals should facilitate B isotope
fractionation, hence the higher δ11B values.
Acknowledgments A part of this study was funded by the Ministry of Energy and Mineral Resources
of Indonesia through PhD scholarship grants number: 2579 K/69/MEM/2010 for B.J.
Purnomo. Thanks to Laura Knigge for the laboratory assistance and to Britta Hinz-
Stolle for an editorial review.
G E O T H E R M A L S Y S T E M S O N T H E I S L A N D O F B A L I , I N D O N E S I A
69
V. Geothermal systems on the island of Bali, Indonesia
Modified from:
Budi Joko Purnomo and Thomas Pichler *
* Geochemistry & Hydrogeology, Department of Geosciences, University of
Bremen, Germany
Submitted to:
Journal of Volcanology and Geothermal Research
G E O T H E R M A L S Y S T E M S O N T H E I S L A N D O F B A L I , I N D O N E S I A
70
Abstract This paper presents an overview of geothermal systems on the island of Bali,
Indonesia by presenting physicochemical data for shallow thermal wells, hot springs,
cold springs and volcanic lakes. A total of 4 locations were sampled and classified
based on their position in either as a volcano-hosted or a fault-hosted geothermal
system. A carbonate host rock for the geothermal reservoirs could not be confirmed,
because the characteristic δ18O shift to the right for carbonate reservoir fluids was not
observed. The HCO3- of the thermal waters was well correlated with Ca2+, Mg2+, Sr2+
and K+ indicating water-rock interaction in the presence of carbonic acid. The (Ca2+ +
Mg2+)/HCO3- molar ratios were approximately 0.4 and K/Mg ratios were typical for
interaction with calc-alkaline magmatic rocks. The B/Cl ratios indicated steam phase
separation for the Bedugul and Banjar geothermal systems. The heavy δ11B of +22.5
‰ and a Cl/B ratio of 820 confirmed seawater input into the Banyuwedang geothermal
system. Comparison of the Si, Na/K, Na/K/Ca and Na/Li geothermometers with actual
reservoir temperature measurements and physicochemical considerations led to the
conclusion that the Na/Li thermometer provided the best results for geothermal
systems on Bali. Using this thermometer, the following reservoir temperatures were
calculated: (1) Penebel (Bedugul) from 235 to 254 °C, (2) Batur 240 °C, (3) Banjar
255 °C and (4) Banyuwedang below 100 °C. The 2H and 18O isotope compositions of
the geothermal fluids indicated meteoric water as a source of all thermal waters
studied.
Keywords: Bali, carbonate, volcanic, host-rock, seawater input, geothermal system,
geothermometer, 2H and 18O isotope
G E O T H E R M A L S Y S T E M S O N T H E I S L A N D O F B A L I , I N D O N E S I A
71
V.1. Introduction The island of Bali, Indonesia is host to several hot springs and geothermal
systems, of which some are of interest for geothermal exploitation. The Bedugul
geothermal field, located near Lake Bratan, was identified to cover an area of
approximately 8 km2 with an estimated annual electric energy potential of 80 MWe for
30 years (Hochstein et al., 2005; Hochstein and Sudarman, 2008; Mulyadi et al.,
2005). However, the development was suspended due to environmental and cultural
concerns. In addition to Bedugul, other geothermal prospects on Bali are the Batur,
Banyuwedang and Banjar geothermal systems.
Bali is dominantly covered by volcanic rocks, overlying the Tertiary carbonate
rocks that outcrop in the southern and western part of the island (Hadiwidjojo et al.,
1998). The Bedugul geothermal field was reported producing brines dominated by
CO2 of approximately 97 wt. % and the reservoir was assumed to be hosted in
carbonate rocks (Mulyadi et al., 2005). To the contrary Geiger (2014) concluded that
the reservoir was in volcanic rocks and that the carbonate basement was leached by
shallow magma, 2 to 7 km depth, in the Batur and Agung volcanoes based on the
thermobarometric study. The dissolution of carbonate rock by magma releases large
amounts of CO2 gas due to the breakdown of CaCO3 into CO2 gas and CaO (Allard,
1983; Chadwick et al., 2007; Deegan et al., 2010; Gertisser and Keller, 2003;
Marziano et al., 2007; Marziano et al., 2009). The rich in CO2 volatile magma
subsequently ascends to the geothermal reservoir and promotes vaporization, which
in turn produces a CO2-rich vapor phase (Lowenstern, 2001).
It is possible that geothermal systems on Bali could be hosted by carbonate
rocks, but CO2 content alone would be insufficient for that conclusion. The carbonate
rocks were determined as the reservoir rock in some volcanic areas in Italy, e.g.,
Vicano-Cimino and Sabatini- Tolfa (Cinti et al., 2011; Cinti et al., 2014). There the
thermal is characterized by a (Ca2+ + Mg2+)/HCO3- molar ratio of ~1 as a result of
calcite and/or dolomite dissolution (Capaccioni et al., 2011; Cinti et al., 2011; Cinti et
al., 2014; Gemici and Filiz, 2001; Goff et al., 1981; Grassa et al., 2006; Levet et al.,
2002; Pasvanoglu and Chandrasekharam, 2011; Pasvanoglu and Gultekin, 2012).
Another characteristic of thermal waters hosted by carbonate rocks could be the
relative 18O isotope enrichment due to the heavier δ18O of the carbonate host rock, as
indicated at Balikesir and Cesme, Turkey; Lanzarote, Spain; and the Salton Sea, USA
(Arana and Panichi, 1974; Craig, 1966; Gemici and Filiz, 2001). Although some
carbonate rock associated geothermal systems also showed depletion of 18O isotope
G E O T H E R M A L S Y S T E M S O N T H E I S L A N D O F B A L I , I N D O N E S I A
72
due to exchange with CO2, for instances at Vicano-Cimino, Sabatini, and Sicily, Italy
(Cinti et al., 2011; Cinti et al., 2014; Grassa et al., 2006).
This paper presents new physicochemical and isotope (18O, 2H and 11B) data for
hot springs and shallow thermal wells on Bali with the objective to investigate the host
rocks of the geothermal systems. Additionally, boron isotopes were applied to identify
seawater input and solute geothermometers to predict the reservoir temperatures.
V.2. Geological setting Bali is a part of the Sunda-Banda volcanic islands arc, which extends for
approximately 4700 km east to west, from the island of Damar to the island of
Sumatera. The arc is caused by the convergence of the Indo-Australian and Eurasia
plates, with a rate of c.a. 6 to 7 cm/a (Hamilton, 1979; Simandjuntak and Barber,
1996). This process caused volcanisms on Bali since the late Tertiary (Hadiwidjojo et
al., 1998; Hamilton, 1979; Van Bemellen, 1949) and produced a vast distribution of
volcanic rocks. The Jembrana volcanic complex occupies the western part of the
island, the Buyat-Bratan-Batur volcanic complex the central part, and the Agung and
Seraya volcanic complexes the eastern parts. Underlying the volcanic rocks are
sedimentary rocks of Tertiary age, which are minimally exposed in the east, south and
west part of the island (Fig. 5.1) (Hadiwidjojo et al., 1998).
The volcanic rocks on Bali are calc-alkaline characterized by a medium
concentrations of K (Nicholls and Whitford, 1983; Whitford et al., 1979). Based on the
thermobarometric results of clinopyroxene and plagioclase at Batur volcano, Geiger
(2014) suggested the existence of a shallow magma in 2 to 4 km depth, associated
with a carbonate sedimentary crust. This is corroborated by InSar satellite data
indicating shallow magma at 2 to 4 km depth (Chaussard and Amelung, 2012) and by
earthquake focal zones at 1.5 to 5 km depth (Hidayati and Sulaeman, 2013).
Volcanoes where carbonate assimilation takes place generally have an explosive
eruption behavior (Deegan et al., 2010).
Explosive eruptions formed two large calderas on Bali, the Batur and the Bratan
calderas (Reubi and Nicholls, 2004; Watanabe et al., 2010; Wheller and Varne, 1986).
Both calderas are geothermal prospects, due to the presence of thermal water in the
shallow subsurface. Although a geothermal reservoir was confirmed beneath the
Bratan lake, (Mulyadi et al., 2005), surface features of geothermal systems, such as
hot springs are virtually absent in the Bratan caldera. Outside of the caldera to the
south, however, several hot springs are present in the Penebel area (Fig. 5.1). On the
G E O T H E R M A L S Y S T E M S O N T H E I S L A N D O F B A L I , I N D O N E S I A
73
northwestern side of the Buyat-Bratan volcano, the Banjar hot spring discharges at
the contact between the Buyat-Bratan-Batur volcanic complex and the Tertiary Asah
Formation. At the western end of the island, the Banyuwedang hot spring is located in
carbonate rocks of the Prapatagung Formation.
Fig. 5.1. Geological map of Bali showing the sampling locations (modified from Hadiwidjojo et al., 1998).
G E O T H E R M A L S Y S T E M S O N T H E I S L A N D O F B A L I , I N D O N E S I A
74
V.3. Results The physicochemical and stable isotopes (18O, 2H and 11B) data of the water
samples from Bali are presented in Table 5.1. The temperatures of thermal waters
ranged from 37.2 to 45.2 °C, while the selected cold waters ranged from 24 to 30.1
°C. The thermal waters had slightly acid to neutral pH, while cold springs neutral and
lake waters were slightly alkaline. The thermal and cold waters had relatively similar
ranges of TDS, Cl-, Na+ and K+ concentrations. Meanwhile the thermal waters had
wider ranges of Ca2+, Mg2+ and HCO3- contents compared to the cold waters. The
Ca2+ concentration of the thermal waters ranged from 51.3 to 211.5 mg/L, Mg2+ from
51.5 to 243.8 mg/L and HCO3- from 31.7 to 2235 mg/L, while for cold waters the Ca2+
and Mg2+ were lower than 100 mg/L and HCO3- ranged from 19.5 to 761.3 mg/L. The
Cl- content of thermal waters ranged from 17.3 to 902.1 mg/L and for cold waters from
under detection limit to 1025.7 mg/L. The Cl-rich cold waters were found in B12 (Batur
Lake) and B8 (Pejarakan), with Cl- contents of 188.6 and 1025.7 mg/L, respectively.
Thermal waters of B1, B2, B3, and B4 had Ca2+, Mg2+ and HCO3- concentrations a
magnitude higher than the other thermal waters. However, B4 differed from B1, B2
and B3 due to its higher SO42- and lower Cl- concentrations. B4 had a SO4
2- content of
111.7 mg/L and a Cl- content of 61.2 mg/L, compared to SO42- of below detection limit
and Cl- ranged from 377 to 443.9 mg/L in B1, B2 and B3. Thermal waters of B6 and
B13 had TDS of <1000 mg/L and Cl- < 20 mg/L, lower than the other thermal waters,
which varied between 1430 to 2600 mg/L and from 61.2 to 902.1 mg/L, respectively.
The two thermal waters probably were more diluted by shallow groundwater. The
thermal water located near the coastline, B7, had the highest TDS and Cl- contents
and the lowest HCO3- concentration.
The thermal waters had δ2H and δ18O values ranging from -42.4 to -33.2 ‰ and
from -6.8 to -5.6 ‰, respectively. This was probably a result of the range in elevation
from close to 0 (seawater) to approximately 1200 m above sea level. The δ2H of cold
springs and a shallow well ranged from 36.9 to -30.0 ‰ and δ18O ranged from -6.0 to -
5.4 ‰. In contrast, cold waters from two freshwater lakes, Batur and Bratan, had
heavier δ2H, ranging from -16.4 to -14.6 ‰, and δ18O, from –2.3 to -1.7 ‰. The
selected thermal waters had a wide range of δ11B compositions, from +1.3 to +22.5
‰. In accordance to the TDS and Cl- contents, the heaviest δ11B value was found in
sample B7, might be due to seawater input.
G E O T H E R M A L S Y S T E M S O N T H E I S L A N D O F B A L I , I N D O N E S I A
75
V.4. Discussion V.4.1. Geochemistry of thermal waters
A geothermal system generally produces three types of hot springs, neutral
chloride, acid sulfate and bicarbonate waters, but mixtures between the individual
groups are common (Hedenquist, 1990; Hochstein and Browne, 2000; Nicholson,
1993; White, 1957). The discharge composition of thermal springs is controlled by two
sets of processes: 1) deep reservoir conditions (deep reservoir = reaction zone
immediately above the heat source), and 2) secondary processes during ascent. In
the deep reservoir, host rock composition, temperature, direct magmatic contributions
and residence time are the controlling factors. During ascent a drop in pressure and
temperature can initiate phase separation and mineral precipitation, causing a
dramatic change in fluid composition. Mixing with other hydrothermal fluids and/or
groundwater is possible at any depth. In near-shore and submarine environments
mixing with seawater cannot be ruled out. The chemical composition of a
hydrothermal fluid, sampled at the surface, generally contains an imprint of its
subsurface history and chemically inert constituents (tracers) provide information
about their source, whereas chemically reactive species (geoindicators) record
physico-chemical changes (Ellis and Mahon, 1977; Giggenbach, 1991; Nicholson,
1993). Classification of thermal waters on Bali using the Cl-SO4-HCO3 ternary
diagram (Chang, 1984; Giggenbach, 1991; Giggenbach, 1997) indicated a
bicarbonate (HCO3-) type for B1, B2, B3, B4, B6 and B13, a mixing type for B9, B10
and B11, and a neutral chloride (Cl-) type for B7 (Fig. 5.2). Neutral chloride waters are
usually thought to represent the deep reservoir fluid, while acid sulfate and
bicarbonate waters form by underground absorption of vapors separated from a
neutral chloride water into cooler ground water. Whether acid sulfate or bicarbonate
waters are formed depends on the gas content of the vapor and redox conditions in
the shallow subsurface (Ellis and Mahon, 1977; Giggenbach, 1997; Hedenquist, 1990;
Henley and Ellis, 1983).
Tabl
e 5.
1. S
ampl
ing
loca
tions
, phy
sico
chem
ical
and
sta
ble
isot
opes
(2 H, 18
O a
nd 11
B) d
ata
of th
erm
al a
nd c
old
wat
ers
on B
ali.
The ρC
O2 w
as
calc
ulat
ed u
sing
the
com
pute
r cod
e P
HR
EEQ
C (P
arkh
urst
and
App
elo,
199
9).
ID
Loca
tion
T pH
O
RP
Con
d.
TDS
Ca
Mg
Na
K C
l H
CO
3 S
O4
Si
Al
Li
B S
r Fe
δ18
O
δ2 H
δ11B
Lo
g
°C
mV
uS/c
m
mg/
l ‰
ρC
O2
Hot
spr
ings
B
1 Y
eh P
anas
1
38.8
6.
5 25
30
16
2238
12
2.4
161.
6 26
3.4
50.1
37
7.0
1466
.4
<dl
76.0
0.
2 1.
3 8.
0 0.
4 <d
l -6
.5
-33
- -0
.5
Pen
ebel
B
2 Y
eh P
anas
2
38.8
6.
6 0
3052
22
72
122.
5 16
4.0
270.
3 51
.2
363.
7 15
25.0
<d
l 76
.0
0.2
1.3
7.6
0.4
<dl
-6.1
-3
6 -
-0.5
P
eneb
el
B3
Yeh
Pan
as 3
42
.6
6.4
-41
3282
24
53
135.
1 16
1.2
309.
4 58
.4
443.
9 15
55.5
<d
l 80
.8
0.2
1.4
9.1
0.4
<dl
- -
10.4
-0
.4
Pen
ebel
B
4 Be
lula
ng
41.8
6.
5 -4
7 34
02
2555
21
1.5
243.
8 23
4.5
67.2
61
.2
2235
.0
111.
7 72
.8
0.2
1.4
4.3
0.8
<dl
-6.8
-4
1 4.
0 -0
.2
Pen
ebel
B
6 A
ngse
ri 45
.2
6.1
1 13
63
943
54.3
81
.7
123.
0 40
.0
16.6
63
4.4
166.
0 97
.3
0.1
0.7
5.4
0.2
0.7
-5.8
-3
3 -
-0.6
P
eneb
el
B7
Ban
yuw
edan
g 44
.6
7.8
-310
34
53
2600
51
.3
51.6
52
6.8
14.8
90
2.1
31.7
20
0.2
11.7
0.
2 1.
2 1.
1 0.
3 <d
l -6
-3
6 22
.5
-3.2
B
13
Ban
jar
37.2
6.
2 -4
1 12
65
873.
6 68
.4
66.2
10
9.2
23.6
17
.3
773.
5 2.
2 73
.1
0.1
0.6
1.9
0.2
1.3
-6.1
-3
7 1.
7 -0
.6
Ther
mal
sha
llow
wel
ls
B9
Toya
Bon
gkah
39
.9
7.5
163
2122
15
23
46.0
75
.8
294.
2 24
.0
159.
3 46
3.6
370.
3 56
.6
0.2
1.3
1.9
0.1
<dl
-6.4
-4
2 1.
3 -1
.8
Bat
ur
B10
Ti
rta H
usad
a 43
.1
7.3
191
2007
14
30
46.6
69
.7
277.
8 22
.8
136.
2 45
8.7
325.
2 62
.0
0.2
1.3
2.0
0.1
<dl
-6
-42
- -1
.6
Bat
ur
B11
To
ya D
evas
ya
40.6
7.
4 16
3 20
55
1478
47
.4
71.1
28
1.6
22.8
14
7.0
488.
0 32
8.7
57.2
0.
3 1.
3 2.
0 0.
1 <d
l -6
.8
-42
- -1
.7
Bat
ur
Lake
sB
12
Bat
ur L
ake
27.2
8.
5 13
5 21
22
1525
31
.8
67.2
32
1.2
26.0
18
8.6
336.
7 48
8.6
1.3
0.3
1.4
1.7
0.0
<dl
-1.7
-1
6.4
- -3
.0
B15
B
rata
n La
ke
24
8.7
195
49.1
2 32
.2
4.6
0.3
<dl
0.8
<dl
19.5
2.
7 0.
0 0.
2 0.
6 0.
4 <d
l <d
l -2
.3
-14.
5 -
-4.4
C
old
sprin
gs a
nd s
hallo
w w
ell
B5
Belu
lang
30
.1
7.0
36
1057
72
7.9
48.9
49
.5
48.6
13
.4
4.3
575.
8 28
.0
27.5
0.
1 0.
6 1.
0 0.
2 <d
l -6
-3
0 -
-1.3
B
8 P
ejar
akan
29
7.
4 93
49
74
3891
40
.0
74.5
87
4.9
68.9
10
25.7
76
1.3
361.
5 31
.0
0.2
1.2
1.1
1.0
<dl
-5.7
-3
7 -
- B
14
Ban
jar
30.9
7.
7 14
3 87
6.2
598.
5 55
.4
41.2
64
.2
15.6
2.
6 56
1.2
4.7
52.3
0.
1 0.
6 1.
1 0.
3 <d
l -5
.4
-31
- -1
.9
was
og CO
2
0.5
0.5
0.4
0.2
0.6
3.2
0.6
1.8
1.6
1.7
3.0
4.4
1.3 -
1.9
G E O T H E R M A L S Y S T E M S O N T H E I S L A N D O F B A L I , I N D O N E S I A
77
Fig. 5.2. Cl-SO4-HCO3 ternary diagram (Giggenbach diagram). Most of the thermal waters were of the bicarbonate type. The samples from Batur (B9, B10 and B11) were a mixing type and B7 a chloride type. The plots of the Pejarakan cold well (B8) and Lake Batur (B12) as mixing water type probably were caused by seawater and thermal water inputs.
The TDS value of 1525 mg/L of B12 was relatively similar to those of the Batur
thermal waters (B9, B10 and B11), probably due to major discharge of thermal water
into the lake. However, it should be noted that B12 was sampled from a site relatively
close to of Batur thermal waters, hence considering the large dimension of the lake of
approximately 6.6 km length and 2.5 km width, the sample probably does not
represent the general chemistry of the lake water. Meanwhile, the higher Cl-
concentration of the shallow well (B8) was likely caused by seawater input due to its
location close to sea.
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Solely based on their geological setting geothermal systems on Bali can be
divided into two groups, volcano-hosted are Batur (B9, B10 and B11) and Penebel
(B1, B2, B3, B4 and B6), while Banyuwedang (B7) and Banjar (B13) are fault-hosted.
Following the interpretation of Purnomo and Pichler (2014) only Banyuwedang (B7)
was considered a truly fault-hosted geothermal system. This thermal water differs
from the others due to its poor of HCO3- content and thus plots as group C, fault-
hosted geothermal system in Figure 5.3. The volcanic-hosted thermal waters are
divided into group A, diluted thermal waters, and group B, originated from the margin
of the ‘primary neutralization’ zone of Giggenbach (1988). In this zone, the elevated
HCO3- content is produced by separation of CO2 and the subsequent reaction with
groundwater. The extent of groundwater mixing governs the Cl- concentration of
thermal waters. The HCO3- content of thermal waters was well correlated (R2 ~ 0.9)
with the Ca2+, Mg2+,Sr2+ and K+ (Fig. 5.4), hence indicated rock dissolution by carbonic
acid. The rocks dissolution was triggered by the formation of weak acid, H2CO3, at
temperatures below 300 °C due to oxidation of CO2 by groundwater (Bischoff and
Rosenbauer, 1996; Giggenbach, 1997; Lowenstern, 2001). The (Ca2+ + Mg2+)/HCO3-
molar ratios of the thermal waters of approximately 0.4 (Fig. 5.5) are lower than the
ratios for thermal waters hosted by carbonate rocks of 1, as reported for the
geothermal systems of Vicano-Cimino, Sabatini-Tolfa and Sicily, Italy; Cesme,
Nevsehir and Terme-Karakurt, Turkey, Bagneres de Bigorre, France; and Jemez
springs, USA (Capaccioni et al., 2011; Cinti et al., 2011; Cinti et al., 2014; Gemici and
Filiz, 2001; Goff et al., 1981; Grassa et al., 2006; Levet et al., 2002; Pasvanoglu and
Chandrasekharam, 2011; Pasvanoglu and Gultekin, 2012). Following this
interpretation a carbonate type host rock for the geothermal systems on Bali could be
ruled out. The thermal waters are more likely hosted by magmatic rocks of calc-
alkaline origin as indicated in the K/Mg vs. Na/K diagram (Fig. 5.6). This type of
magma is generally produced by volcanoes located in the middle of a volcanic island
arc, in between tholeiitic and high-K calc-alkaline magmatic regions (Whitford et al.,
1979).
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Fig. 5.3. HCO3 vs. Cl diagram. Groups A and B were volcano-hosted thermal waters originated from the ‘primary neutralization’ zone, where group B was less diluted by groundwater compared to group A. Group C was fault-hosted thermal water, hence underwent insignificant magmatic fluid input.
The final pH of the thermal water was governed by addition of CO2, hence
H2CO3 formation, and subsequent water-rock interaction. While addition of CO2 lowers
the pH, water-rock interaction causes an increase due to the comsuption of H+. The
effect of CO2 addition in Penebel group (B1, B2, B3, B4 and B6) and Banjar (B13)
dominated water-rock interaction and thus kept a slightly acid pH (Fig. 5.7a).
Meanwhile, the CO2 supply in Batur group (B9, B10 and B11) and thus decreasing pH
was neutralized by water-rock interaction. However, the lower ρCO2 and higher pH of
Batur group compared to the other thermal waters also could be caused by calcite
precipitation, which also releases CO2, indicated by the saturation condition with
respect to calcite (Fig. 5.7b).
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Fig. 5.4. a) Ca vs. HCO3, b) Mg vs. HCO3, c) K vs. HCO3 and d) Sr vs. HCO3 diagrams. The well correlations of HCO3
- with Ca2+, Mg2+, K+ and Sr2+ indicated rock dissolution by carbonic acid.
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Fig. 5.5. The (Ca2+ + Mg2+)/HCO3- molar ratios of thermal waters were approximately
0.4, an indication of non-carbonate host-rocks. For comparison (Ca2+ + Mg2+)/HCO3-
molar ratios of the carbonate-hosted thermal waters plotted using data from Cinti et al. (2014), Capaccioni et al. (2011) and Levet et al. (2002).
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Fig. 5.6. K/Mg vs. Na/K diagram indicates dissolution of calc-alkaline volcanic rocks of Batur volcano. Data of calc-alkaline volcanic rocks of Batur are from Reubi and Nicholls (2004) and high-K volcanic rocks of Muria volcano from Whitford et al. (1979).
Fig. 5.7. a) log ρCO2 vs. pH and b) SI calcite vs. pH diagrams. Thermal waters with high ρCO2 had lower pH, a typical of CO2 fed thermal waters. The water pH probably was also buffered by calcite precipitation.
V.4.2. Phase separation and seawater input The Cl/B ratio of thermal waters can be used to identify water-rock interaction,
steam separation and seawater input in the subsurface (Arnorsson and Andresdottir,
1995; Purnomo and Pichler, 2014; Valentino and Stanzione, 2003). The Cl vs. B
diagram of thermal waters from Bali indicates reaction with andesitic rock for the Yeh
Panas group (B1, B2 and B3) and the Batur group (B9, B10 and B11) of thermal
waters. Either B depletion or seawater input is indicated for Banyuwedang (B7) and
steam phase separation for Belulang (B4), Angseri (B6) and Banjar (B13) (Fig. 5.8).
The last process elevates the B/Cl ratio of the thermal water by scavenging B into the
steam phase, thus relatively increasing the Cl content of the residual water (Arnorsson
and Andresdottir, 1995; Truesdell et al., 1989). That steam phase separation occurred
for B4 and B6 was confirmed due to the existence of two phases, liquid and vapor,
zone in the liquid-dominated reservoir of the Bedugul geothermal field (Hochstein et
al., 2005).
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The distinct δ11B composition of seawater of +39.61 ‰ (Foster et al., 2010) is
heavier than other fluid sources and thus can be used to investigate seawater input or
B adsorption onto minerals to explain the B/Cl ratio observed for Banyuwedang (B7).
This method was successfully applied for some geothermal systems on Java island,
Indonesia (Parangtritis and Krakal), Iceland (the Reykjanes and Svartsengi) and three
areas in Japan (the Izu-Bonin arc, Kusatsu-Shirane area, and Kagoshima) (Aggarwal
and Palmer, 1995; Aggarwal et al., 2000; Kakihana et al., 1987; Millot et al., 2009;
Musashi et al., 1988; Oi et al., 1993; Purnomo et al., 2015). Adsorption of B by
minerals also increases the δ11B of the thermal water due to the preferentially
fractionation of 10B into the solid phase (Palmer et al., 1987; Schwarcz et al., 1969;
Xiao et al., 2013). The magnitude, however, is lower than what can be observed due
to seawater input. Banyuwedang (B7) had a δ11B composition of +22.5 ‰ and a Cl/B
ratio of 806; therefore, it plots close to the mixing line between seawater and thermal
water in the δ11B vs. B/Cl diagram (Fig. 5.9), which would confirm seawater input. The
thermal water shift away from the seawater point in the B vs. Cl diagram was caused
by groundwater mixing that lowered its B and Cl- concentrations.
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Fig. 5.8. Cl vs. B diagram illustrates four processes in the sub surface, i.e., a steam phase separation for B4, B6 and B13; an andesitic rock leaching for B1 to B3 and B9 to B11; and either a B depletion or seawater input for B7.
Fig. 5.9. The plot of Banyuwedang close to the mixing line of seawater-thermal water in the δ11B vs. B/Cl diagram confirms seawater input.
V.4.3. Geothermometry
The reservoir temperatures of geothermal systems on Bali were calculated
using solute geothermometers, SiO2, Na-K-Ca, Na-K and Na-Li (Table 5.2) (Fournier,
1977; Fournier, 1979; Fournier and Truesdell, 1973; Kharaka and Mariner, 1989).
These geothermometers record their temperature-dependent equilibrium reactions,
hence application of multiple geothermometers can be used to evaluate secondary
processes during thermal water ascent from the reservoir to the surface. These
processes include shallow water dilution, conductive cooling, adiabatic cooling,
mineral precipitation, water-rock interaction and re-equlibration (Fournier, 1977;
Kaasalainen and Stefánsson, 2012). Such an evaluation has been successfully
applied, for instances, on Java, Indonesia and Ambitle island, Papua New Guinea
(Pichler et al., 1999; Purnomo and Pichler, 2014).
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The silica geothermometer calculated lower reservoir temperatures compared to
the Na-K, Na-K-Ca and Na-Li geothermometers, which is common for samples taken
at the surface from hot springs, rather than directly from the geothermal reservoir.
That geothermometer predicted reservoir temperatures ranging from 44 to 136 °C,
while those predicted by the Na/K thermometer ranged from 257 to 773 °C, those
predicted by the Na/K/Ca thermometer from 130 to 236 °C and those predicted by the
Na/Li thermometer from 162 to 254 °C (Table 5.2). The calculation of silica
geothermometry is based on absolute silica content, hence sensitive to boiling,
precipitation and dilution (e.g. Nicholson, 1993). This deficiency can be overcome by
calculating the silica parent using the silica mixing model of Fournier (1977). However,
this method could not be used on Bali because the thermal waters in a given
geothermal system had relatively similar temperatures and silica contents, for
example the Penebel group (B1, B2, B3, B4 and B6). Meanwhile, the Na/K
temperatures are likely overestimations caused by competition of Ca2+, Na+ and K+
during ion exchange (Nicholson, 1993). The use of the Na/Li geothermometer resulted
in reservoir temperatures ranging from 235 to 254 °C for the Penebel thermal waters
(B1, B2, B3, B4 and B6), which were relatively similar to actual reservoir temperatures
at 1800 m below ground of the nearby Bedugul geothermal field of 243 °C (Mulyadi et
al., 2005). This indicates the applicability of the Na/Li geothermometer as has been
proposed by, e.g., Fouillac and Michard (1981). Based on this, the reservoir
temperatures of the other geothermal systems on Bali, with an exception of
Banyuwedang (B7), were predicted using Na/Li geothermometer. Therefore, the
reservoir temperature of the Batur geothermal system was approximately 240 °C and
Banjar was 255 °C. However, due to the input of seawater in Banyuwedang (B7),
there a geothermometer based on Na+ content is unreliable. The silica
geothermometer predicted a temperature of 44 °C for B7, similar to the discharge
temperature and hence a likely underestimation. Therefore, the reservoir temperature
of B7 could not be reliably calculated, but probably is lower than 100 °C, a
temperature similar to most of the fault-hosted geothermal system on Java (Purnomo
and Pichler, 2014).
Calcite precipitation during thermal water ascent was predicted by comparing
the result of Na/Li and Na/K/Ca geothemometers. Precipitation of calcite during
thermal water ascent reduces the Ca2+ concentration and thus should result in lower
calculated Na/K/Ca temperatures. Calculated temperatures ranged from 209 to 236
°C for the Penebel group of thermal waters and thus were similar to those calculated
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the Na/Li geothermometer, which would indicate insignificant calcite precipitation
during fluid ascent. In contrast, the difference between 170 °C (Na/K/Ca) and 240 °C
(Na/Li) for the Batur group of thermal waters indicates that calcite precipitated during
ascent. This is corroborated by the lower ρCO2 and higher pH compared to the
Penebel thermal waters (Fig. 5.7).
Table 5.2. Calculated reservoir temperatures using solute geothermometers.
Location Sample Geothermometers (°C) ID Silica Na/K/Ca Na/K Na/Li
(Truesdell, 1975)
(Fournier and Truesdell, 1973)
(Fournier, 1979)
(Kharaka et al., 1982)
Penebel
B1 122 209 568 244 B2 122 209 567 242 B3 126 211 566 235 B4 120 227 688 254 B6 136 236 733 254
Banyuwedang B7 44 130 257 190
Batur B9 108 172 395 236 B10 122 171 396 240 B11 108 171 394 239
Banjar B13 120 205 602 255
V.4.4. Oxygen and hydrogen isotope considerations The deuterium and oxygen isotopic composition of thermal waters has been
successfully applied to investigate the fluid origin, i.e., meteoric, marine or magmatic,
mixing and physicochemical processes, such as, water-rock interaction and water-
CO2 isotope exchange (Arnason, 1977; Cinti et al., 2014; Craig et al., 1956; Gemici
and Filiz, 2001; Giggenbach et al., 1983; Pichler, 2005; Purnomo and Pichler, 2014).
The δ2H and δ18O composition of water is generally a good indicator of its origin.
Hydrothermal fluids normally plot to the right of the LMWL due to exchange of 18O
during water-rock interaction (Craig, 1966) or due to subsurface mixing with an
andesitic water, as defined by (Giggenbach, 1992). In the δ2H vs. δ18O diagram all
thermal waters, except the two lake waters, plot close to the local meteoric water line
(LMWL) and weighted mean annual value for precipitation in the region, indicating
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local rainwater as the ultimate fluid source (Fig. 5.10). With the exception of sample
B10 the thermal waters from Bali did not shift to the right of the LMWL indicating
neither water-rock isotope exchange or mixing with an andesitic water. Neither
process, however, can be completely ruled out because an initial 18O-shift to the right
may have been later reversed by a subsequent isotope exchange between CO2 and
H2O. Such isotope shifts were observed in several CO2-rich aquifers and
hydrothermal waters (Chiodini et al., 2000; Cinti et al., 2011; Cinti et al., 2014; Grassa
et al., 2006; Vuataz and Goff, 1986). The Tirta Husada hot spring (B10) from the Batur
group plots slightly right shifted from Toya Devasya (B11), potentially signaling some
water-rock interaction. However, the TDS of B10 was relatively similar to B11 and B9
(Fig. 5.11). Although shallow magma is present at the Batur volcano (Geiger, 2014),
an isotope isotope enrichment by magmatic input was not observed. The absence of a
pronounced δ18O shift to heavier values also points towards a non-carbonate
reservoir. Due to water-rock interaction in carbonate reservoirs heavier δ18O values
are generally observed at such locations (Pasvanoglu and Chandrasekharam, 2011).
Fig. 5.10. δ2H vs. δ18O diagram with the local meteoric water line (LMWL) from Wandowo et al. (2001), GMWL from Craig (1961) and the mean annual value for
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precipitation in Jakarta from IAEA (2008). The thermal waters were of meteoric water origin. B10 is slightly horizontal right shifted from B11, probably was caused by water-rock interaction. The heavy water isotope of Lake Bratan and Lake Batur were caused by evaporation.
Fig. 5.11. TDS vs. δ18O diagram shows relatively similar TDS for the Batur thermal waters, B9, B10 and B11, hence indicating an insignificant 18O enrichment due to water-rock interaction.
The elevated δ2H and δ18O compositions of Lake Batur (B12) and Lake Bratan
(B15), which were studied in detail by Varekamp and Kreulen (2000), were a result of
evaporation in a cold lake. B12 was slightly enriched in δ18O compared to B15, likely
due to the significant input of thermal waters into the lake that elevated its
temperature. A lake with a higher temperature would produce a heavier δ18O due to
evaporation (Gonfiantini, 1986; Varekamp and Kreulen, 2000).
V.5. Conclusions Two types of geothermal systems are present on Bali, the fault-hosted for
Banyuwedang geothermal system and the volcano-hosted Penebel, Batur and Banjar
geothermal systems. Contrary to assumptions by others we conclude that, although
Bali is underlain by a carbonate basement, the geothermal reservoir rocks are of calc-
alkaline magmatic origin. Steam phase separation occurred in Penebel and Banjar
geothermal systems, while seawater input was confirmed for the fault-hosted
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geothermal system of Banyuwedang. Comparison of several geothermometers with
actual reservoir temperature measurements and physicochemical considerations led
to the conclusion that the Na/Li thermometer provided the best results for geothermal
systems on Bali. Using this thermometer, the following reservoir temperatures were
calculated: (1) Penebel (Bedugul) from 235 to 254 °C, (2) Batur 240 °C, (3) Banjar
255 °C and (4) Banyuwedang below 100 °C. The deuterium and oxygen isotopic
composition indicated the thermal water were of meteoric water origin without
significant of water-rock interaction and/or mixing with an andesitic magmatic fluid. Acknowledgments
B.J. Purnomo likes to thank the Ministry of Energy and Mineral Resources of
Indonesia for the PhD scholarship grants number: 2579 K/69/MEM/2010. Thanks to
Chen-Feng You for the boron isotope measurement, to Laura Knigge for the
laboratory assistance, to Ketut Suardana for the help during fieldwork and to Britta
Hinz-Stolle for an editorial review.
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VI. Conclusions and outlook
VI.1. Conclusions
Geothermal systems in the Sunda volcanic island arc are mostly a volcano-
hosted system. In this geological setting, the geothermal systems are considered
mainly influenced by Quaternary volcanic activities. However, a closer look at the
geothermal systems on Java showed that a fault-hosted geothermal system could
exist exclusively, without any influence from the Quaternary magma. Specific
conclusions coming from this study are:
1. The geothermal systems on Java could be classified into two types, volcano-
hosted and fault-hosted. However, the classification could not be solely based on
the location, either in a volcano or fault, magmatic fluids input has to be taken into
account. Geothermal systems hosted by fault distributed in the Quaternary volcanic
belt were supplied by magmatic fluids, thus could not be classified as a fault-
hosted geothermal system.
2. The presence of magmatic fluids input in the volcano-hosted geothermal system
elevated the HCO3-, Mg/Na ratios and stable isotope (2H and 18O). These
characteristics were not identified in the fault-hosted geothermal systems.
3. In the fault-hosted geothermal systems deep circulated groundwater is conductive
heated. Quaternary magma fluids input are absent due to the relatively
impermeable Tertiary volcanic rocks. However, the Quaternary magma could be
the heat source for the fault-hosted geothermal system, as indicated in the Cilayu
and Cisolok geothermal systems.
4. Another heat source for the fault-hosted geothermal systems located in the Tertiary
volcanic belt was considered a deep-seated magma, like has been identified
elsewhere, e.g., the Alpine fault, New Zealand (Allis and Shi, 1995; Shi et al.,
1996). The plates subduction in the south of Java has uplifted and thinned the
crust of the sourthern part of the island, hence deep circulated groundwater can
extract a deep-seated magma.
5. The boron isotope composition of thermal waters further distinguished volcano-
hosted and fault-hosted geothermal systems. The fast ascent and absence of
C O N C L U S I O N S A N D O U T L O O K
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magmatic fluids input kept a light δ11B value of the fault-hosted geothermal systems
and conversely, slow ascent and magmatic fluids input, which is favorable for
minerals precipitation, promoted δ11B enrichment in the volcano-hosted geothermal
systems.
6. The contrasting processes between acid chloride and acid sulfate crater lakes
produced different B isotope signatures. The former had light δ11B values
representing direct magmatic fluids supply from the subsurface. In contrast, the
latter were δ11B enriched by vapor phase separation in the subsurface, followed by
evaporation and B adsorption into clay minerals after discharge.
7. The study of geothermal systems on Bali identified two types of geothermal
systems, a fault-hosted for Banyuwedang and a volcano-hosted for Penebel
(Bedugul), Batur and Banjar. The thermal waters had (Ca2+ + Mg2+)/HCO3- ratios of
approximately 0.4 and stable (2H and 18O) isotope typical of meteoric water origin,
hence ruled out a carbonate host rock type. The host rocks are volcanic of calc-
alkaline magma, as indicated by the K/Mg. Vapor phase separation was identified
occurring in Penebel (Bedugul) and Banjar geothermal systems. The fault-hosted
geothermal system of Banyuwedang is influenced by seawater input.
VI.2. Outlook This study reveals a complex geothermal systems in the Sunda volcanic island
arc. The fault-hosted geothermal systems are not necessarily influenced by the
Quaternary volcanic activities. Since every location has unique geological
characteristics, the methods used in this study could be applied in other volcanic
islands arcs, e.g., Japan, Philippines and New Zealand. The studies will benefit for
better understanding the geothermal systems in such a geological setting, hence the
model can be established. The absence of Quaternary magmatic fluid should be
corroborated in the next studies by incorporating sulfur (34S of SO42-) and carbon (13C
of CO2) isotopes to trace the origin. Sulfur in a geothermal water can be sourced from
1) seawater, 2) magmatic gases (H2S and SO2) and 3) oxidation of sulfide minerals
(Pichler, 2005). 13C isotope of CO2 was applied to trace the contribution of mantle
degassing and rock leaching, e.g., in Tutum Bay, Papua New Guinea by Pichler
(2005) and the Sabatini volcanic area, Italy by Cinti et al. (2011). The geochemical
study can be supported by details geological mapping and geophysical measurement
to configure the subsurface condition in the boundary of the Quaternary volcanic belt
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and Tertiary volcanic belt. Although the Tertiary volcanic rocks are relatively
impermeable, the presence of fractures potentially circulating fluids.
Concerning on the boron isotope signature of the two contrast thermal crater
lakes, acid sulfate and acid chloride, this study shows an overview of their different
boron isotope signatures and provides the possible explanations. More samples
representing both acid crater lakes from different location would corroborate this
finding. The distinct characteristic of the two crater lakes, the Kawah putih, rich in B
and light in δ11B values, and Kawah Sikidang, rich in B and heavy in δ11B values,
deserve for a more detail study. A time series of sampling should be done to
understand the fluctuation of δ11B isotope value. The behavior of B isotope after
discharge can be determined by performing experiments of B isotope fractionation
during evaporation and clay adsorption for low pH water with high and low Cl-
concentrations, representing acid sulfate and acid choride, respectively. To date the
effect of pH on B isotope fractionation during water-clay interaction is only known for
higher pH (> 5), e.g., Yingkai and Lan (2001).
The study on Bali provides evidences of a non carbonate reservoir type, thus
the rich of CO2 in the geothermal brines likely was resulted by CO2-rich magma
volatile due to a carbonate assimilation, as was proposed by Geiger (2014). 13C
isotope of CO2 has to be incorporated in the upcoming studies. Access to the
geothermal brines of the Bedugul geothermal field could be used to examine its
contrasting physicochemical properties compared to the geothermal brines from Java
(Bogie et al., 2008; Mulyadi et al., 2005; Purnomo and Pichler, 2014). The indication
of significant thermal water supply into the Lake Batur needs to be confirmed by a
systematic sampling. The sampling also can be applied to map the physicochemical
characteristic distribution, leading to an identification of thermal water sources. The
significant thermal water input might be originated from underwater thermal water
discharge, hence a survey of underwater hot spot is needed.
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A C K N O W L E D G E M E N T S
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Acknowledgements
This PhD was made possible by funding from the Ministry of Energy and Mineral
Resources of Republic Indonesia and University of Bremen.
I would like to first thank my supervisor, Prof. Dr. Thomas Pichler for the opportunity to
do PhD, something beyond my imagination before you invited me. I learned the
passion of doing fieldwork as a geoscientist from the experience working along with
you during sampling campaign across Java in 2012. Words will not be enough to
express my gratitude for your advices, supports and critical discussions during my 4
years doing PhD.
I thank to Prof. Dr. Peter LaFemina for the agreement to review and examine my PhD
thesis. Thanks to Prof. Dr. Chen-Feng You for the boron isotope analysis. The fourth
chapter of this thesis would not be possible without your contributions.
Special thank to Laura Knigge, Christine Albert and Christian Breuer for the help
during laboratory analyses, also Britta Hinz-Stolle for your much effort and time on the
editorial review of my manuscripts. Wisanukorn Poonchai is thanked for your German
to English translation during our PHREEQC summer course. I also thank to Dr. Kay
Hamer and Dr. Lars-Eric Heimburger for the fruitful advices, discussions and inputs.
Pak Ketut, Maman and Seto are thanked for the assistance during sampling campaign
on Java and Bali.
My fellow PhD students in the Geochemistry and Hydrogeology group, Wu Debo, Qi
Pengfei and Ali Mozafarri, for the support, sharing, discussion and encouragement.
Special for my office mates, Debo and Pengfei, thanks for always keep a warm
atmosphere in our beloved office.
I thank to my fellow ‘Cator’-ers (Indonesian PhD students in Bremen): Riza Setiawan,
Mba Asty, Mba Hesty, Teh Rima, Kang Ayi, Condro, Yusup, Haliq, Mba Enab, Mba
Meity, and Rovi. Our togetherness kept my motivation and spirit on finishing this
study.
A C K N O W L E D G E M E N T S
96
Finally, my beloved wife, Novi, and daughter, Inca, deserve very special thank. I
dedicate this for you both. Thanks to my big family, my father and mother (in heaven),
my 9 brothers and sisters, Mas Kentut, Mba Pon, Mba Iten, Mba Sri, Mas Tris, Mba
Nik, Mas Cipto, Mba Yani and Ulin. I could not achieve this level without all of you. My
parents in law are thanked for taking care my small family during I stay in Germany.
R E F E R E N C E S
97
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