Journal of Volcanology and Geothermal Research 331 (2017) 26–52
Contents lists available at ScienceDirect
Journal of Volcanology and Geothermal Research
journal homepage: www.elsevier.com/locate/jvolgeores
Geochemical study of the Sakalol-Harralol geothermal field (Republic of
Djibouti): Evidences of a low enthalpy aquifer between Manda-Inakir
and Asal rift settings
Mohamed Osman Awaleh a,⁎, Tiziano Boschetti b, Youssouf Djibril Soubaneh c, Paul Baudron d,e,f,
Ali Dirir Kawalieh a, Omar Assowe Dabar a, Moussa Mahdi Ahmed a, Samaleh Idriss Ahmed a,
Mohamed Ahmed Daoud a, Nima Moussa Egueh a, Jalludin Mohamed a
a
Centre d'Etudes et de Recherches de Djibouti (CERD), Route de l’aéroport, B.P. 486, Djibouti, Djibouti
Department of Physics and Earth Sciences “Macedonio Melloni”, University of Parma, Parco Area delle Scienze 157/a, 43124 Parma, Italy
c
Département de biologie, chimie et géographie, Université du Québec à Rimouski, 310, Allée des Ursulines, Rimouski, QC G5L 3A1, Canada
d
Département des génies civil, géologique et des mines, Polytechnique Montréal, C.P. 6079. succ. Centre-Ville, Montréal, QC H3C 3A7, Canada.
e
UMR G-EAU, BP 5095, 34196 Montpellier Cedex 5, France
f
GEOTOP Research Center, Montréal, Canada
b
a r t i c l e
i n f o
Article history:
Received 7 September 2016
Received in revised form 11 November 2016
Accepted 11 November 2016
Available online 14 November 2016
Keywords:
Afar
Djibouti
Hydrogeochemistry
Isotopes
Geothermometry
Sakalol
a b s t r a c t
Eighty-six sodium bicarbonate to sodium chloride hot springs and four water wells in the Tadjourah Region of
Djibouti were investigated for major, minor (B, Br, F, Sr, Li) chemistry and isotope composition of water and
dissolved components (87Sr/86Sr, 11B/10B, 13C/12C and 14C of DIC, 34S/32S and 18O/16O of sulfate). The deep saline
Na-Cl reservoir at 143 °C shows affinity with the shallow geothermal water from the “active” Asal rift. Asal water
is a diluted and recycled seawater component with the major cation composition obliterated by equilibration
with Stratoid basalt. Locally, the deep reservoir is differentiated in term of recharge, and re-equilibration with
rocks and mixing. In particular, two spring groups reveal contributions from evaporites typical of the “passive”
graben setting of the Afar. A model on 34S/32S and 18O/16O demonstrates the isotope imprint of magmatic SO2 disproportionation on dissolved and solid sulfate, whose values probably persists in a sedimentary environment
without trace of seawater. On the other hand a seawater signature, modified by mixing and secondary fractionation effects, is partially maintained according to the boron isotope composition (up to +27.4‰). Temperature
estimation in low-enthalpy geothermal reservoirs is notoriously difficult, especially where mixing with fluids
of differing genesis and/or conduction cooling take place. From a geothermometric point of view, the multimethod approach followed in this study (up-to-date theoretical and thermodynamic equations, ad-hoc silica
geothermometers inferred from local rocks, checking of the results on a 18Oαsulfate-water vs. temperature diagram)
provides some insights and perspectives on how to tackle the problem.
© 2016 Elsevier B.V. All rights reserved.
1. Introduction
As in other rifting zones, the activity of the East African Rift System
corresponds to large seismic, tectonic and volcanic activities (Barberi
et al., 1975; Mlynarski and Zlotnicki, 2001). Such a particular
geodynamical situation gives to the Republic of Djibouti a remarkable
position for the development of geothermal energy. Indeed, the
Republic of Djibouti features numerous hot springs, fumaroles and
hydrothermal alteration mainly distributed in the western part of the
country and along the Gulf of Tadjourah ridge (Fig. 1A). However, few
studies have been completed on areas with geothermal activities in
⁎ Corresponding author.
E-mail address: awaleh@gmail.com (M.O. Awaleh).
http://dx.doi.org/10.1016/j.jvolgeores.2016.11.008
0377-0273/© 2016 Elsevier B.V. All rights reserved.
the Republic of Djibouti for estimating their geothermal potential
(Sanjuan et al., 1990; D'Amore et al., 1998; Awaleh et al., 2015a,b).
In the Sakalol and Harralol horst-and-graben structure extending towards south down to north the Asal Rift, there are a dozen thermal
springs. The Sakalol-Harralol geothermal field (~ 200 km2) is one of
the largest geothermal fields of the Republic of Djibouti, which is located
in the north of the East Africa Rift. However, due to remote and poorly
accessible sites, the geochemistry of 6 thermal springs from the
Sakalol-Harralol geothermal field (SHGF) only was reported in the
early seventies (BRGM, 1970). Systematic geochemical sampling and
analysis of the numerous hydrothermal springs in SHGF was needed
for a better understanding of the regional geothermal regime of this
wide area located between two active spreading segments of Asal and
Manda-Inakir (Fig. 1).
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
The main purpose of the present study is to characterize the hydrothermal activity of the SHGF. Detailed geochemical investigations have
been carried out for the first time on the cold groundwaters (well and
boreholes waters) and almost all thermal waters (86 thermal springs)
from the Sakalol-Harralol area. Environmental and radiogenic isotopes
34
11
87
2−
(δ2H-H2O, δ18O-H2O, δ18O-SO2−
Sr/86Sr, δ13C, 14C)
4 , δ S-SO4 , δ B,
as well as major and minor ions chemistry are applied to assess the geochemical evolution and groundwater residence times of the thermal
waters from the SHGF. The results of this investigation are compared
with previous published data on the hot waters from neighboring graben and rift systems.
Boron isotopes (δ11B) have over the last two decades frequently
been used to elucidate groundwater salinization and pollution problems
(e.g. Vengosh et al., 1994, 1998, 1999). To the best of our knowledge, in
no study δ11B value neither for groundwater nor for thermal waters in
the East Africa Rift system has been reported. In the present study the
first results of boron isotopes in thermal waters from the East Africa
Rift system are reported.
The goals of this study are: (1) to classify the thermal water composition data into genetic groups, (2) to characterize the main geochemical processes that explain the thermal waters geochemistry and to
understand their geochemical evolution, (3) to estimate the temperature in the SHGF geothermal reservoir through chemical and isotopic
geothermometry and (4) to estimate the eventual scaling processes of
the SHGF geothermal waters. The results obtained allow for the understanding of the behavior of this system on a regional scale and are useful
for planning future management of the Sakalol-Harralol geothermal
system.
2. Geological and tectonic setting
The SHGF is located within the northwestern portion of the
Ghoubbet-Asal, which forms an accretionary emerged rift segment penetrating the Afar depression (Varet, 1975). Similar geological features are
the Manda Inakir, the Manda Hararo and the Erta Ale. The Afar depression
is a triple junction characterized by thinned continental crust, where
three rift systems meet: the Red Sea, the Gulf of Aden and the East
African Rift (Gaulier and Huchon, 1991; Manighetti et al., 1998) (Fig. 1A).
The most complete Afar synrift volcanic succession is well exposed
in the Republic of Djibouti. The oldest units are located in the AliSabieh region. They consists of a mafic intrusion (20–28 Ma, Le Gall
et al., 2010) core mantled by an outward-dipping envelope, including
a Jurassic-Cretaceous sedimentary cover overlain by (i) initial mafic effusive sequences, and (ii) younger acidic series of the 15–11 Ma old
Mablas Fm (Zumbo et al., 1995). These latter are onlapped by the younger Dalha (8.6–3.8 Ma) and Somali (7.2–3.0 Ma) basalts (Chessex et al.,
1975) whose floor depressed the area and the Djibouti plain respectively (Fig. 1A). The youngest volcanic units, related to the recent downfaulting and rifting process comprise Stratoid trap-like basaltic series
b3 Ma, covering most parts of the Afar depression (Varet, 1975;
Zumbo et al., 1995), the Tadjourah Gulf Basalts (2.8–1.0 Ma) associated
with the initiation of Tadjourah rifted zone (Manighetti et al., 1998;
Daoud et al., 2011) and Asal-Ghoubbet and Manda Inakir volcanic at
around 900 ka (Manighetti et al., 1998; Audin et al., 2001) are assumed
to represent the axial part rift system (Vellutini, 1990) (Fig. 1A).
The stratigraphy of the Sakalol-Harralol area comprises the extensive Plio-Pleistocene (2.3–1.0 Ma) Stratoid basalts (Gasse et al., 1987)
which rest unconformably over highly weathered basalts of the Dalha
series (Fig. 1B). In addition, the sedimentary infill consists of alluvial,
salty silt and hydrothermal deposits (Fig. 1B). The hydrothermal calcite
edifices, several meters high, are located in the vicinity of still active
thermal springs, in veneer crowning basaltic escarpment. Their construction began during the Asal high stand (ca 6800 yr. B.P) and stable
isotope contents in hydrothermal concretions indicate higher evaporation rates than a previous phase, when a connection existed between
Asal and Alol (Gasse and Fontes, 1989). The ancient lake had probably
27
no perennial surface, received an inlet from Lake Asal and fed by direct
precipitation surface runoff from the escarpments walls, and finally by
groundwater.
The restricted spatial distribution of active hydrothermal manifestations only along the hanging wall of the normal faults shows that existing
fault structures provide key pathways for the upflow of the geothermal
fluids (Kerrich, 1986; Norton and Knapp, 1977) (Fig. 1B). Hot springs occurrence seems occupying a specific area along fault zones. To the north,
the hydrothermal manifestations are located on the high angle steeply
dipping fault, bounding the northern margin of the sakalol intrabasin.
Hot springs are clustered in the central part of the fault where displacement is maximal according to the classical fault growth models prediction (Cowie and Scholz, 1992). To the south, the hot springs are
situated at the terminations (fault-tips) of the individual segmented
faults, limiting both Harralol intrabasin margins (Fig. 1B). Fault overlaps
zone separating the two Harralol-Haralé intrabasin appears to be a favorable site for the development of hot springs (Fig. 1B). Fault-tips and faults
overlaps are sites of elevated stress causing active fracturing and continual re-opening of fluid-flow conduits, permitting long-lived hydrothermal flow (Hutchison et al., 2015; Kaya, 2015).
The general morphostructural framework of the studied geothermal
zone comprises three adjacent structural units with similar morphological features. It consists of a succession of three elongated half-graben
structures facing NE and their corresponding intra-basins which are
from, south to the north, Harralol, Haralé and Sakalol (Figs. 1B and 2).
Individual intra-basins are generally less than 3-km-wide, and they
are bounded by N140°E striking normal faults, dipping to NE and orthogonal with respect to the inferred N40°E regional tectonic extension
(Gaulier and Huchon, 1991).
The dominant N140°E fault set is composed of linear, highly segmented structures showing at their extremity curved and overlapping
map traces, N 15 km in length (Figs. 1B and 2). Thus, the fault network
along the Alol geothermal field displays steeply dipping attitudes and
shows southerly tilted stratoid basaltic blocs (Fig. 2). Very few antithetic
fault structures exist, for example, that intersects the northern margin of
Sakalol intra-basin, facing SW and causing the vertical downthrow of
the Stratoid Basalts towards the sub-rift axis (Fig. 2). It should be
noted that these faults and corresponding tilted blocks were developed
on the Makarassou westerly facing monocline flexure that lowers the
Stratoid surface down to the Asal rift floor (Tapponnier et al., 1990; Le
Gall et al., 2011). The highly elevated (~1650 m) and segmented relief,
occurring further south in the prolongation of the Asal fault system
corresponds to the footwall of the Doubié master fault (Fig. 2). It presents an overall convex-shaped morphology resulting from an active
tectonic uplift (Doubre et al., 2007) (Fig. 2). This fault exhuming a
large part of the older Dalha series shows a low-angle dip. These observed geometrical features characterize the Doubié fault as a detachment fault.
It is commonly observed that systems containing several subparallel faults that dip and tilt the same way are kinematically coupled
with basal shear zone (Jackson et al., 1988). Following the wide-scale
“conjugate passive margins” concept applied to the entire SE Afar extensional domain, Geoffroy et al. (2014) interpreted recent/active pattern
rift kinematics in the Asal/Alol rifted zone in terms of a crustal extension
in which the Doubié fault acted as detachment that merges into the updomed part of the basal shear plane, beneath the Asal. Consequently, the
Makarassou monocline flexure is considered a post-magmatic roll-over
structure developed in the hanging of the Doubié detachment fault
(Fig. 2).
3. Material and methods
3.1. Sampling and analytical methods
A total of eighty six thermal springs, two well waters and two borehole waters were sampled in April 2015. Temperature (± 0.1 °C), pH
28
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
A
B
Fig. 1. (A): A schematic geological map of the Republic of Djibouti (SE Afar Rift) and hydrothermal activity of the Republic of Djibouti. (B): A schematic geological map of the SakalolHarralol geothermal field (C1: Cluster 1 thermal waters; C2: Cluster 2 thermal waters; C3: Cluster 3 thermal waters; C4: Cluster 4 thermal waters; B: Borehole water; WW: Well waters).
In the inset: schematic map of the Afar Depression with the location of Djibouti (black rectangle).
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
29
Makarassou
monoc
line
flex
ure
ASAL
1670 m Co
alo
l
detachment
gy
lt
ié fau
Doub
rra
ol
ral
é
Sa
k
Tilted b
locs
Ha
Ha
ph o l o
d mor
ul
t
a
Doubié
f
e
-shap
ex
nv
-150 m
~ 4km
Fig. 2. 3D Landsat view of the Doubié detachment fault and tilted blocks developed in the lower part of the Makarassou flexure.
(±0.01 unit) and electrical conductivity (±1 μS/cm), were measured
on site using the following portable instruments: 1) temperature CheckTemp(Hanna); 2) pH - pH 610 (EutechInstruments); and 3) conductivity - COND 610 (Eutech Instruments). The instruments were calibrated in the field.
Water samples were collected into polyethylene containers after filtration through 0.45 μm membrane filters. All samples used for determination of cations were acidified after collection through addition of
Suprapure® HNO3 (Merck) to bring the pH below 2.
Analysis of anions and major cations was carried out with a Dionex
ICS 3000 Ion Chromatograph using analytical and quality assurance procedures for geothermal water chemistry, following Pang and
Armannsson (2006).
Analysis of trace elements (B, Sr) was conducted using an Ultima 2
(Horiba Jobin Yvon) Inductively Couple Plasma Atomic Emission Spectrometer (ICP-AES). National Institute of Standards and Technology
traceable commercial standards were used for quality control. These
standards were analyzed to within ±3% of the known values for trace
elements. For the analysis of aqueous SiO2, the water samples were diluted tenfold using deionized water to prevent SiO2 polymerization.
SiO2 concentrations were determined by colorimetry and analyzed
using a Jenway 6300 spectrophotometer, while HCO3 was analyzed for
by titration with 0.1 M HCl following Gran method. The charge balance
between anions and cations ((Σ[Anions] - Σ[Cations])/(Σ[Anions] +
Σ[Cations])), concentrations in meql− 1, was assessed and analyses
were accepted for deviations of less than 3%.
Additional samples of untreated waters were collected into 50 ml
glass bottles (Quorpak) for analysis for stable isotopes of the water molecule (δ2H-H2O and δ18O-H2O). Water samples filtered through 0.45 μm
membrane filters were collected into 1000 ml Nalgene bottles for
carbon-14, and into 250 ml plastic bottles for carbon-13 determinations
of dissolved inorganic carbon (13C-DIC). The isotope ratios of hydrogen
and oxygen were analyzed for at the GEOTOP laboratory at the
Université du Québec à Montreal (Canada) using a Micromass Isoprime
isotope ratio mass spectrometer and converted into per mil delta values
(δ‰) versus the Vienna Standard Mean Ocean Water (V-SMOW) standard following δ (‰) = [(Rsample/Rstandard) - 1] × 103, where R is the isotopic ratio of interest (2H/1H or 18O/16O). The average precision, based
on multiple analyses of various samples and laboratory standards was
±0.1‰ for δ18O-H2O and ±0.8‰ for δ2H-H2O. The 13C ratios were determined using a mass spectrometer (Micromass, Isoprime model
with triple universal collectors) at the GEOTOP laboratory (Canada).
13
C contents are reported using the conventional δ (‰) notation as a
deviation from the V-PDB standard, and the error for δ13C is 0.05‰
(mean error obtained from replicate analyses).
The 14C activities were measured by accelerator mass spectrometry (HVE 3 MV Tandetron accelerator) at the A.E. Lalonde Laboratory
at the University of Ottawa (Canada). CO2 for 14C analysis was extracted from water samples in a vacuum and converted to graphite
by reducing it with excess hydrogen gas in the presence of a
preconditioned (reduced) Fe powder catalyst. The 14C results were
reported as percent Modern Carbon (pMC) with an average 1σ
error of ± 0.3 pMC.
Water samples for 87Sr/86Sr analyses were filtered through 0.45 μm
membrane filters and acidified to pH b 2 with Suprapure® HNO3
(Merck) before being collected into 250 ml polyethylene containers.
Sr-isotopes were determined in a VG sector 54 in dynamic mode at
the GEOTOP laboratory (Canada). 87Sr/86Sr ratios were normalized to
87
Sr/88Sr = 0.1194. Repeated analyses of the NIST-987 standard yielded
values of 0.710294 (± 0.000022, 2σ reproducibility). Furthermore,
these results were corrected to the accepted value for NIST-987 of
0.710248.
Samples for sulfur and oxygen isotope analyses were collected at the
outlet of columns using 250 ml pre-acid-washed plastic perplex bottles.
Cd-acetate was already added to the bottles (5% v/v) prior to sample
collection (to fix sulfur as CdS) and then the aliquot was filtered through
a 0.2 μm nitrocellulose filter before the chemical determination of residual sulfate. Then, dissolved sulfate was precipitated as BaSO4 at pH b 4
(in order to remove HCO3 − and CO2−
species) by adding a BaCl2 solu3
tion. The isotope analysis of BaSO4 was carried out using an IsoChrom
Continuous Flow Stable Isotope Ratio Mass Spectrometer coupled to a
Carlo Erba Elemental Analyzer at the Environmental Isotope Laboratory
at the University of Waterloo. Sulfate-isotope ratios are reported in the
usual δ-scale in ‰ with reference to V-CDT (Vienna Canyon Diablo Troilite) and V-SMOW (Vienna Standard Mean Ocean Water). Sulfateisotope ratios (δ34S(SO4) and δ18O(SO4)) were determined with a precision of ± 0.3‰ vs. V-CDT for δ34S(SO4) and ± 0.5‰ vs. V-SMOW for
δ18O(SO4).
Water samples for boron isotope analyses were collected into 500 ml
polyethylene containers after filtration through 0.45 μm membrane filters. The B isotope composition was analyzed using a TIMS VG 336 at the
Tetra Tech Laboratory (USA). The 11B data were reported in δ per mil
(‰) relative to the standard of NIST SRM 951. The average value for
the NIST SRM 951 standards analyzed with each batch of samples is normalized to the accepted isotopic ratio; this correction is applied to the
measurement for each sample in a given batch analyzed in the mass
30
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
spectrometer. Subsequently, the standard deviation is computed and
reported as the long-term precision of the isotopic measurement to 1σ.
3.2. Data elaboration
The computer program R (2016) was used to conduct statistical
analysis (Hierarchical Cluster Analysis and Principal Components Analysis – HCA and PCA). The HCA was performed using a combination of
Ward's linkage method (Ward, 1963), adopting the Euclidean distance
as measure of dissimilarity. The PCA was implemented using the
FactoMineR (Factor analysis and data Mining with R) package
(Husson et al., 2009). Calculation of mineral saturation indices,
logP(CO2) and activity of dissolved component were performed using
PHREEQCI, version 3 and thermo.com.v8.dat thermodynamic dataset
(Parkhurst and Appelo, 2013). This latter and Thermoddem (release
2014; Blanc et al., 2012) datasets were used to draw mineral stability
fields in activity diagrams by The Geochemist's Workbench, version 7
(Bethke and Yeakel, 2008). Finally, full chemical composition and carbon isotope data were elaborated together by PHREEQCI and the iso.dat
dataset to obtain δ13C(CO2) from δ13C(DIC), whereas Netpath XL was
used to calculate and correct groundwater age using 14C data
(Parkhurst and Charlton, 2008).
4. Results and discussion
Field data (temperature, pH, electrical conductivity, total dissolved
solids, sampling locations) major and minor elements and hydrochemical
types of all water samples are listed in the supporting information file and
summarized in Table 1. Isotope results are in Table 2.
4.1. Statistical analysis
HCA and PCA have commonly been used as a quantitative and independent approach for groundwater classification (Meng and Maynard,
2001; Swanson et al., 2001; Güler et al., 2002). HCA is generally used
to classify observations or variables, in order to define more or less homogeneous groups and emphasize their genetic relations (Davis, 1986).
In other words, HCA produces the most distinctive groups where each
member within the group is more similar to its fellow members than
to any member outside the group (Güler et al., 2002).
In this study, 13 variables (EC, Ca, Mg, Na, K, HCO3, Cl, SO4, NO3, F, Br,
B, and SiO2) in 90 water samples were analyzed using Q-mode HCA and
PCA.
The result of the HCA is presented as a dendrogram (Fig. S1, see
supporting information). The HCA allowed us to distinguish four clusters (C1–C4) of thermal water in the Sakalol–Harralol geothermal field
(Fig. S1): Cluster 1 (less mineralized thermal waters with EC =
1500–4740 μS/cm); Cluster 2 (thermal waters with EC =
4215–9516 μS/cm); Cluster 3 (thermal waters with EC =
7804–12412 μS/cm), and Cluster 4 (thermal water with EC =
22660 μS/cm). The robustness of the HCA results was verified by relating the statistically defined clusters to their geographical locations. In
fact, the four clusters had shown a well-distinguished distribution,
with good correspondence between spatial locations and the statistical
groups (Fig. 1B). Samples that belong to the same group are located in
close proximity to one another, suggesting different steps of the
evolutive processes along different flow paths (Güler and Thyne,
2004). In other terms, it was reported that waters which fall into a statistical group may have similar residence times, a similar recharge history, and identical flow paths or reservoirs (Swanson et al., 2001; Güler
and Thyne, 2004).
PCA of water chemical variables from the SHGF produced two components (Table S1, see supporting information), accounting for 74.18%
of the total variance of the dataset. The principal component loadings,
as well as their respective explained variances are presented in
Table S1. Loadings, that represent the importance of the variables for
the components are in bold for values greater than 0.7 (Table S1).
Most of the variance is contained in the PC1 (63.12%) which is associated with the variables EC, Ca, Na, K, HCO3, Cl, SO4, Br and B, with
squared regression coefficients of 0.93, 0.69, 0.98, 0.87, − 0.77, 0.98,
0.86, 0.95 and 0.81 respectively. Therefore, PC1 can be used to account
for a Sakalol–Harralol thermal waters salinization processes. PC2
(11.06%) is mainly related to F and SiO2 with squared regression coefficients of 0.84 and 0.70 respectively. Therefore PC2 is conceived to represent mainly a dimension that describes mineralization of thermal
waters from the SHGF by silicate dissolution.
Table 1
Mean chemical and physicochemical parameters of the four statistically determined water groups in the Sakalol-Harralol geothermal field. See the supporting information file for specific
values in the hot spring and local non-thermal groundwater samples.
Parameter
EC
pH
T
TDS
Ca
Mg
Na
K
Li
NH4
HCO3
Cl
SO4
NO3
NO2
F
Br
SiO2
Sr
B
Unit
(μS/cm)
at sampling temperature
(°C)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
(mg/l)
Cluster 1
Cluster 2
Cluster 3
*Cluster 4
Na–Cl–HCO3-SO4
Na–Cl
Na–Cl
Na–Cl
N = 34
N = 31
N = 20
N=1
Mean
St.dev.
Mean
St.dev.
Mean
St.dev.
Mean
St.dev.
1928
8.22
45.9
1043
10.8
2.82
380
14.7
0.058
1.05
197
366
164
15.1
0.040
4.95
1.21
92.7
0.062
0.792
654
0.38
4.9
365
5.8
2.28
132
3.4
0.060
2.40
53
193
36
4.3
0.039
0.86
0.57
10.2
0.017
0.172
6607
8.12
42.2
3537
45.0
6.37
1258
40.5
0.096
0.369
120
1766
273
9.15
0.24
2.62
5.37
95.8
0.28
1.44
1509
0.32
7.6
745
37.2
2.11
249
12.0
0.102
0.395
42
405
52
4.85
0.40
1.51
1.56
9.2
0.14
0.42
9778
7.91
45.8
5289
112
13.7
1844
50.1
0.119
1.98
68
2676
493
8.10
0.071
4.72
8.71
107
0.64
2.67
1294
0.33
5.9
596
42
8.6
285
15.2
0.110
5.15
23
409
91
7.13
0.052
1.08
1.41
11
0.31
0.92
22660
7.16
77.7
13265
1020
12.9
3653
119
1.03
0.53
3.16
7073
399
2.92
0.030
2.06
27.1
123
15.9
3.12
0.01
1.6
72
3.4
152
5.7
0.12
0.40
1.38
396
18
0.40
0.014
0.20
2.1
1
0.02
N: number of sampled springs (2015).
⁎ : Cluster 4 is composed of only one spring (Galiceela). Data in these columns refer to two samplings carried out twice (2015 and 2016).
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
31
Table 2
Isotope data for the Sakalol-Harralol geothermal field.
Cluster #
Sample
#
Sample name
δ18O(H2O)
(‰ vs
V-SMOW)
δ2H(H2O)
(‰ vs
V-SMOW)
δ13C(DIC)
(‰ vs
V-PDB)
*δ13C(CO2)
(‰ vs
V-PDB)
14
C
(pmc)
δ11B
(‰ vs
SRM951)
δ34S(SO4)
(‰ vs
V-CDT)
δ18O(SO4)
(‰ vs
SMOW)
87
1
1
4
6
9
10
12
16
19
21
23
25
30
31
32
33
35
36
37
39
43
46
47
49
51
53
60
48
61
62
65
66
67
70
71
72
74
75
76
77
78
79
80
81
82
84
86
88
90
douffalou 1
douffalou 4
douffalou 6
ciddayta 3
halmalé 1
dir dara 1
konadara 2
konadara 5
konadara 7
konadara 9
konadara 11
konadara 16
As dara 1
As dara 2
kouroubta 1
Dora borehole
Balho borehole
hagande well 2
hamadaba 1
hamadaba 8
hamadaba 11
hamadaba 12
Dor Asi 1
Dor Asi 3
ina citta 3
hamadaba 6
hamadaba 14
hagandé 3
hagandé 6
As bodara 3
As bodara 1
As bodara 2
badacii 1
badacii 2
badacii 3
Ina citta 1
ina citta 2
hamadaba 4
birtibodeh
loubag dara
ali celah
Aden Dara
houmad nouh 1
houmad nouh 2
houmad nouh 4
houmad nouh 6
houmad nouh 8
galiceela
−3.90
−3.87
−4.07
−3.88
−3.89
−3.88
−4.04
−3.79
−3.60
−3.22
−0.27
−2.74
−3.27
−3.28
−2.18
−1.88
−3.49
−3.1
−2.22
−2.12
−2.94
−3.00
−2.87
−2.25
−2.29
−3.12
−2.34
−2.29
−1.96
−2.19
−2.12
−2.15
−2.26
−1.72
−28.1
−27.6
−28.6
−26.4
−28.0
−27.4
−26.8
−25.3
−24.6
−20.7
1.20
−18.0
−24.4
−21.5
−15.3
−12.9
−24.2
−20.3
−17.7
−14.2
−22.6
−23.5
−21.0
−20.2
−18.5
−23.6
−19.0
−16.6
−17.3
−19.3
−18.0
−19.9
−19.2
−12.9
−6.89
−6.08
−6.54
−6.68
−6.74
−12.97
−5.13
−5.50
−7.27
−6.95
−6.97
−4.57
−2.38
−3.12
−9.75
−13.35
−12.69
−12.97
−13.42
−13.28
−18.86
−12.26
−12.63
−14.30
−12.60
−13.11
−10.74
−8.86
−9.81
−14.42
18.10
19.04
20.69
21.37
20.73
29.27
33.07
35.63
32.29
21.21
20.79
23.54
24.47
23.31
25.71
12.8
18.5
20.5
13.5
24.2
11.4
19.1
18.7
21.0
27.4
8.16
7.60
8.75
9.67
8.73
9.88
8.54
9.82
8.63
10.69
6.60
11.46
12.29
12.54
13.2
12.71
9.74
14.13
13.66
8.79
13.44
13.10
14.61
14.28
11.49
11.75
11.87
12.42
11.43
11.86
11.82
11.92
11.08
11.25
8.79
11.68
10.38
12.01
11.55
11.24
9.96
11.42
11.00
10.16
11.25
11.6
11.88
11.39
0.7057
0.70530
0.7053
0.7061
0.7056
0.7055
0.7057
0.7056
0.7054
0.7053
0.7063
0.7041
0.706
0.7055
0.7058
0.7055
0.7055
0.7044
0.7044
0.7046
0.7058
0.7057
0.7037
well. bore.
2
3
4
Sr/87Sr
⁎ : Calculated by PHREEQCI and iso.data thermodynamic-isotopic dataset (Parkhurst and Appelo, 2013).
The water sample observations of the two principal component
axes are plotted in Fig. S2 (supporting informations). In score plot
PC1 vs. PC2 (Fig. S2), the water samples could be subdivided into
four distinct clusters, which are quite similar to the groupings obtained by HCA.
4.2. Main hydrochemical features and classification
Hydrothermal activity in the SHGF is characterized by thermal springs
that discharge from faults generally aligned NW-SE (Fig. 1B). The temperatures of the thermal water samples at the SHGF ranged from 38 to
77.7 °C, with an average of 45.1 °C (Table 1). Geothermal waters are moderately alkaline with an average pH of 8.11 (pH = 7.02–8.95).
Classification of the waters samples in Table 1 was made according to the principles of IAH (IAH, International Association of
Hydrogeologists, 1979). Total equivalents of cations and anions
were separately accepted as 100% and ions with more than 20%
(meq/l) were taken into consideration in this classification. Moreover, chemical composition of the waters was plotted in anion and
cation ternary diagrams on weight basis (Fig. 3), the former with
the typical waters fields detected in geothermal areas as described
by Giggenbach (1991a), and Langelier-Ludwig in equivalent basis
(e.g. Boschetti, 2011; Fig. 4). According to statistical analysis, the
chemical composition of the waters described in terms of relative
concentrations of the main anions and cations allows the discrimination of the following four groups of waters (Table 1, Fig. 3):
(1) Cluster 1 (C1): less mineralized thermal waters from the SHGF
(TDS = 715–2455 mg.l−1) are mostly of the Na–Cl–HCO3–SO4
type. The high alkalinity of these waters is typical of the Rift
groundwaters where high rate of carbon dioxide outgassing
32
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
ma
Obock
eC
tu r
lw
a te
Abhe
rs
T.A.D.D.
rip
pe
vo
lc a
nic
wa
ter
s
Minkileh
/sh
ral
HCO3
ow
all
SO4
he
Cl
rs
te
wa
(dug well)
#35
steam-heated waters
Mg
Ca
this study
well-borehole
Na+K
Cluster 1
Sakalol
Cluster 2
Cluster 3
Cluster 4
Obock
Abhe
literature
geothermal well
T.A.D.D.
Oued-Kalou
Asal
hot spring
Lake Asal
geothermal well
Mg
Hanlé
Ca
borehole
springs
Na+K
seawater
Obock
Abhe
T.A.D.D.
Fig. 3. Anions (A) and cations (B1 and B2) ternary diagrams (data in mg/l). Diamonds are hot spring waters and borehole waters from Sakalol (this study). In (A), peripheral/shallow,
volcanic and mature fields are from Giggenbach (1988, 1991a). In both, gray fields depict other geothermal waters from Afar and Djbouti: Obock (Awaleh et al., 2015a); Abhe (Awaleh
et al., 2015b, Houssein et al., 2013; JICA, 2014); T.A.D.D. (Tendaho-Allallobeda-Dobi-Dubti graben system, Ali, 2005; UNDP, United National Development Program, 1973; Panichi,
1995; D'Amore et al., 1997). Water samples from Asal (Bosch et al., 1977; Correia et al., 1985; D'Amore et al., 1998; Sanjuan et al., 1990; JICA, 2014) and Hanlé (Aquater, 1986; BFGUR,
1999; Houssein, 2010; JICA, 2014) are also shown for comparison. From this latter area, Minkileh spring and Hanlé 1, Hanlé 2 geothermal wells have been analyzed in this study
(squares with thick border).
from mantle and which enhances rock water interaction at shallow depth (e.g. Kebede, 2013):
(4) Cluster 4 (C4): saline thermal water (TDS = 13265 mg.l−1) are
of the Na–Cl type. It is represented by the highest temperature
sample from Galiceela (#90).
CO2 þ H2 O þ Na; K−silicates→HCO3 þ Na; K þ H−silicates
(2) Cluster 2 (C2): moderately mineralized thermal water (TDS =
2199–4790 mg.l−1) are of the Na–Cl type.
(3) Cluster 3 (C3): brackish thermal water (TDS = 4334–
6946 mg.l−1) are of the Na–Cl type.
Notwithstanding the main Na-Cl composition, waters from the SHGF
are rich in sulfate that is typical of geothermal waters in rift settings in
the region (D'Amore et al., 1997; D'Amore et al., 1998; BFGUR, 1999;
Awaleh et al., 2015b). The concentration increases proportionately
with salinity, therefore distinct sulfate concentrations are revealed
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
33
(SO4 + Cl)
T.A.D.D.
50
#35
#75
Obock
45
40
waters interacting with
sedimentary formations
Abhe
#90
Na-HCO3 waters
sensu stricto
35
(Ca + Mg)
(Na + K)
30
25
20
A
0
0
0
5
10
15
20
25
30
35
40
45
50
(HCO3)
(SO4 + Cl)
50
Minkileh
Na-HCO3 waters
sensu stricto
45
Dobi
Obock
waters interacting with
sedimentary formations
40
Abhe
35
(Ca + Mg)
(Na + K)
30
25
20
15
10
5
B
0
0
0
5
10
15
20
25
30
35
40
45
50
(HCO3)
well/borehole
Cluster 1
Sakalol
Cluster 2
Cluster 3
geothermal well
Asal
hot spring
Lake Asal
geothermal well
Hanlé
borehole
seawater
hot-warm-cold springs
Cluster 4
Fig. 4. Langelier-Ludwig diagram (data in meq/l). Sample symbols as in Fig. 1. In (A), waters from Sakalol (this study) and Asal are plotted. In (B), waters from Hanlé. In both, a bricked field
was depicted by waters interacting with a sedimentary formation in the Hanlé field (Aquater, 1986; Houssein, 2010).
between water groups: C1 = 164 ± 36 mg/l, C2 = 273 ± 52 mg/l, C3 =
486 ± 93 mg/l, except Galiceela spring (C4) whose sulfate concentration was not significantly different from C3 group: 399 ± 18 mg/l. In
the literature, the relatively high quantity of sulfate in the East Africa
Rift System has been linked to the circulation of relatively deep waters
within a thick sedimentary surface sequence (D'Amore et al., 1997;
34
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
Awaleh et al., 2015b). However, the anions ternary diagram (Fig. 3)
shows that most of the C1 water group evolves from the HCO3, shallow
(ground)water corner, towards the “volcanic waters” side at a constant
SO4/Cl ratio of 0.36 ± 0.07. It is interesting to note that this value is similar to that revealed in the waters from Hanlé-Gaggade (Hanlé 1 and 2
geothermal well: 0.27 ± 0.02, N = 2, BFGUR, 1999; Minkileh spring:
0.37 ± 0.02, N = 3, JICA, 2014, this work) and Dobi (0.24 ± 0.12,
N = 19; UNDP, United National Development Program, 1973, Panichi,
1995; Ali, 2005, 2011) grabens, which have a widely sedimentary
cover (e.g. Tesfaye et al., 2008).
Therefore, the origin of sulfate could be related both to a sulfate of
deep hydrothermal source, as occurring in Asal area (D'Amore et al.,
1998; Fouillac et al., 1989), and to sulfates of sedimentary origin.
In the anions ternary diagram, C2 waters plot close to the geothermal waters of the Ethiopian Afar (T.A.D.D. waters in Fig. 3: Tendaho,
Allallobeda, Dobi, Dubti) and show a vertical trend up to C4 water,
whereas C3 waters resemble the geothermal waters of Lac Abhe. Despite the increase of chloride concentrations in the water groups, clusters 1 to 3 fall out of the “mature Cl field” similarly to other Afar Rift
geothermal waters. C4 waters represent the exception, plotting very
close to Asal hot springs samples. The presence of seawater near to
the chloride apex of the diagram can lead to misleading interpretations
of a direct involvement of this end-member in the SHGF circulation
(Fig. 3). In fact, it should be noted that a cation diagram shows distinctly
that Asal springs plot between seawater and local mature geothermal
water, according to their mixed marine origins (D'Amore et al., 1998;
Fouillac et al., 1989; Sanjuan et al., 1990), whereas most of the SHGF waters group in the 90% sector of alkali vertex, plot like other Afar hot waters of mainly meteoric origin (T.A.D.D. field). In the same cations
diagram, C4 waters approach Asal geothermal waters plotting like
those of Abhe hot springs, but this latter similarity does not appear in
the anions plot. This means that the cations acquisitions processes are
by water-rock interaction with conditions similar to the Lake Abhe geothermal waters, whereas the anion concentrations are controlled by the
deep gases, mainly SO2/H2S and CO2, and high the amount of chloride
from the Asal geothermal system.
Finally, the Langelier-Ludwig plots combine the informations from
both ternary diagrams (Fig. 4). Here it is evident that all waters from
the Sakalol-Harralol are mainly of alkali-chloride type, except that
from the Dora borehole (#35) which has a Na-bicarbonate signature.
The straight line departing from the Na-bicarbonate corner separates
clusters 1–3 from C4, where the formers are more similar to T.A.D.D.
geothermal waters, whereas the latter to Asal-Abhe geothermal waters
(Fig. 4a).
Most of the waters from Hanlé show a distribution similar to that of
C1–C3 from Sakalol, but with more scattering probably due to a contribution from sedimentary cover. In particular, waters from geothermal
wells and boreholes (Hanlé 1-2 and R1-R4, respectively; BFGUR,
1999) drilled and sampled in the 1980-1990s plots near that of the
Dobi hot springs, whereas Hanlé geothermal well waters sampled in
2009 (Houssein, 2010) and 2016 (this study) show an increasing contribution of dissolved sulfate or meteoric/sedimentary waters of bicarbonate composition, respectively.
4.3. Water isotope composition
The oxygen δ18O(H2O) and hydrogen δ2H(H2O) ratios are useful
tracers for determining the origin of groundwater and are widely used
in studying natural water circulation and groundwater movement.
The stable isotope compositions of the SHGF waters are shown in
Table 2. In Fig. 5, all the data are plotted along with: the Global Meteoric
Water Line (GMWL) δ2H(H2O) = 8 × δ18O(H2O) + 10 defined by Craig
(1961); the Local Meteoric Water Line (LMWL) as defined by Fontes
et al. (1980), with the same slope of the GWML but null deuterium excess (d = 0) (Fig. 5, Awaleh et al., 2015a,b); other water isotope data of
Djibouti from literature (Aquater, 1986; Awaleh et al., 2015a,b; BFGUR,
1999; Bosch et al., 1977; Correia et al., 1985; D’Amore et al., 1998; Darling and Talbot, 1991; Fontes et al., 1979, 1980; Fouillac et al., 1989;
Houssein, 2010; JICA, 2014; Panichi, 1995; Schoell and Faber, 1976;
Verhagen et al., 1991) and other data from geothermal systems of the
Main Ethiopian Rift (M.E.R.) and Ethiopian Afar (Ali, 2005; Craig et al.,
1977; D'Amore et al., 1997; Gonfiantini et al., 1973; IAEA, 2007; Panichi,
1995). The typical and most evident geothermal 18O-shift was detected
in the Danakil depression, in particular in Dallol hot springs waters
(Gonfiantini et al., 1973), but a mixing trend towards basaltic water
has recently been suggested after analysis of fumarolic condensates in
the area (Franzson et al., 2015; Taran and Zelenski, 2015). In the case
of Djbouti hot springs, meteoric and geothermal waters plot very close
to the LMWL (Awaleh et al., 2015a,b). The SHGF samples from this
study confirm this tendency, being mainly grouped near to the LMWL
at δ2H b− 10‰. This “deuterium divide” is in accord with a deep recharge for Abhe and Hanlé geothermal waters (Awaleh et al., 2015b).
Exception is represented by the Dora borehole water (sample # 35)
which is more shifted towards enriched 2H values similarly to the
Wadi-recharged aquifer (Awaleh et al., 2015b), the Awash river and diluted springs from Hanlé (e.g. Aquater, 1986). Interpretation of this
sample may be ambiguous because the sample falls within the standard
deviation of geothermal water from Asal geothermal wells (Fig. 5).
However, its δ2H N−10‰ and an alignment on a straight line parallel
to meteoric water lines along with the Balho borehole (#36) and Cluster
1 samples suggests a shallow meteoric recharge, compatible with nonevaporated monsoon rains which have generated heavy floods
(Awaleh et al., 2015b). Another evidence from the water isotope plot
is that Cluster 1 and Cluster 2 waters from this study show a distribution
comparable with the “Hallol-Harallol” springs water data published
previously (Fig. 5; Fontes et al., 1980, where Hallol-Harallol = SakalolHarralol), with Cluster 1 thermal waters showing the most depleted
δ2H and δ18O values in comparison with any other hot spring waters
from the study area (Fig. 5, Table 2). In particular, the distribution of
Cluster 1 samples is between Hanlé, Danab and Abhe geothermal waters. The depleted isotope composition of Cluster 1 thermal waters
could probably be due to a recharge from meteoric water from a higher
altitude than the other SHGF thermal water groups. This recharged
groundwater is probably from Moussa Ali Mountain (altitude
~ 2021 m) groundwater which flow towards Sakalol-Harralol depressions (BGR, 1982).
Dor Asi and Hagande waters of the Cluster 2 hot springs (#49–51
and #61–62, respectively) and Balho borehole (#36) groundwater
(depth ~300 m, T ~62 °C) show similar stable isotopes values as Abhe
hot springs, as well as Hanlé 2 geothermal water (Fig. 5; Awaleh et al.,
2015b), whereas Hanlé cold boreholes waters were mainly recharged
from the regional aquifer (Aquater, 1986; BFGUR, 1999). Therefore,
these thermal waters from the SHGF should originated from a geothermal reservoir fed by meteoric water mainly from the regional aquifer, as
is the case with the Lake Abhe geothermal field (Awaleh et al., 2015b).
All but one Cluster 3 hot springs plot very close to the LMWL, with
values slightly shifted to the right side interpretable as an 18Oenriched composition due to water-rock interaction at high temperature and/or low water/rock ratio (Fig. 5). However, mixing with basaltic
water could not be excluded, considering the similarity between the
mixing path depicted with that showed by literature data for Abhe hot
springs (Awaleh et al., 2015b; Fontes et al., 1980).
Finally, the C4, Galiceela (#90) hot spring water in the study area has
a very similar stable isotopes composition to the Korili hot spring of meteoric origin from the Asal geothermal field (JICA, 2014; Fontes et al.,
1989), located in the south of the study area (Fig. 5). This indicates
that probably the same meteoric groundwater supplies the geothermal
reservoirs from which Galiceela (#90, C4 thermal water) and Korili/
Oued Kalou hot springs rise up from the SHGF and Asal geothermal
fields, respectively. In other terms, this meteoric groundwater transits
from the Asal local rift to the Sakalol-Harralol depression without any
significant change in the isotope content. It has been reported that
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
35
this study
well-borehole
20
10
Lake
Asal
#35
0
Cluster 1
Cluster 3
Cluster 4
literature
Schoell & Faber 1976
Fontes et al. 1979, 1980
Adam 1984
Verhagen et al. 1991
Dallol
(hot springs)
-10
-20
mean ± s.d.
GMWL
(d=10)
Dallol
(fum.cond.)
geothermal well
Erta Ale
(fum.cond.)
-40
Asal?
Afar and M.E.R.
geotermal fluids
T.A.D.D.
samples
pooling
-30
2
Sakalol
Cluster 2
Asal
M.E.R.
δ H(H2O) (‰ vs V-SMOW)
n
tio
ra
po )
a
ev = 5
(S
r
ve
Ri
sh
a
Aw seawater
A
Asal
hot spring
LMWL
(d=0)
-50
-60
Basaltic
Water
-70
-80
-6
-4
-2
0
2
4
6
8
10
18
δ O(H2O) (‰ vs V-SMOW)
-10
B
-12
#90
-14
literature
-16
-18
borehole
-20
-22
Hanlé
deep geothermal well
(Hanlé 2; BFGUR 1999)
-24
-26
Danab
-28
Abhe
-30
2
δ H(H2O) (‰ vs V-SMOW)
spring
#36
-32
Abhe springs
(Fontes et al. 1980)
-34
-36
-38
-40
Asal springs?
(Fontes et al. 1980)
-42
-5
-4
-3
-2
-1
0
1
18
δ O(H2O) (‰ vs V-SMOW)
Fig. 5. Hydrogen vs. oxygen stable isotope composition. Global (GMWL, Craig, 1961) and local (LMWL, Fontes et al., 1980) meteoric water lines are shown. Sample symbols as Fig. 3, but
open diamonds: SHGF from literature. See the main text for a complete list of the references relating to the samples showed. In (A) and (B), arrows depict hypothetical mixings with
“basaltic water” (Allard, 1983; Kyser and O'Neil, 1984; Giggenbach, 1991b). In (A), dashed line represents best fit of the Awash River data (Ayenew et al., 2008; Panichi, 1995; JICA,
2015; Yitbarek et al., 2012). The line with a slope of about 5, i.e. representing evaporation effect, been traced on the CK sample (Canal Sakalol, Fontes et al., 1980) and extended up to
LMWL. In (B), a detail of Sakalol hot springs waters is shown.
seawater supplies most of the hot springs located northward from the
Lake Asal, while meteoric groundwater supplies Korili and Oued Kalou
hot springs located southwestward from Lake Asal (e.g. Fontes et al.,
1989; Sanjuan et al., 1990). Therefore, since there is no isotopic evidence of seawater intrusion from the Asal local rift to the SakalolHarralol basin, it is more likely that there is an underground connection
36
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
only between the South of the Asal local rift and the Sakalol-Harralol depression. The supposed communication of Lake Asal with the northern
depressions of Sakalol and Harralol by paleochannels (Gasse and
Fontes, 1989) is therefore confirmed by our isotope data.
It is of interest to note that the Galiceela (#90) hot spring in the
study area is about 50 km far from Korili hot spring (Lake Asal area).
Water from geothermal well Asal 1, drilled in the Lake Asal geothermal
field, has different stable isotopes ratios and chloride concentration than
Cl (mg/l)
18
δ H(H 2O) (‰ vs V-SMOW)
10
seawater
5
0
A1 well
-5
Korili
-10
-15
-20
5
-25
#35
0
-30
-5
-10
well-borehole
Cluster 1
Cluster 2
Oued-Kalou
Minkileh
this study
Sakalol
Cluster 3
#90
-15
T.A.D.D.
Cluster 4
-20
literature
Schoell & Faber 1976
Verhagen et al. 1991
Agna
Abhe
-25
Danab
geothermal well
Asal
hot spring
-30
0
spring
Hanlé
1000
2000
3000
4000
5000
6000
7000
borehole
geothermal well
(Hanlé 2; BFGUR 1999)
Dobi
waters with Cl > 1 g/l
Cl (mg/l)
18
δ O(H2O) (‰ vs V-SMOW)
2
seawater
1
0
-1
A1 well
Korili
-2
-3
-4
0.5
T.A.D.D.
0.0
#35
-5
-0.5
-1.0
Oued-Kalou
Minkileh
-1.5
#90
-2.0
-2.5
Abhe
-3.0
Agna
-3.5
Danab
-4.0
-4.5
0
1000
2000
3000
4000
5000
6000
7000
Fig. 6. Water isotope vs. chloride concentration. Symbols as in Fig. 5. Continuous line: high chloride concentration – low isotope values limit of the samples from Sakalol. Dashed line:
fittings between A1 well (Fouillac et al., 1989) and sample C4 from this study (#90, Galiceela). Dotted and dashed-dotted lines represent hypothetical fittings of the hot springs from
Hanlé and Sakalol with hot springs interacting with alluvial/evaporitic deposits (Oued-Kalou, Bosch et al., 1977; Dobi, Ali, 2005), respectively. T.A.D.D.: Tendaho-Allallobeda-DobiDubti graben system. See main text for details and complete references list of the samples from literature.
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
37
for the application of solute geothermometers (Fig. 7). This diagram
is widely used in high enthalpy (T N 150 °C) and convergent margin
systems. However, in low enthalpy (T b 150 °C) and basaltic rift systems, the Giggenbach's square plot (Giggenbach, 1988; Giggenbach,
1995) may also give useful information related to secondary processes that may affect water and rocks (Fig. 8). In the ternary diagram (Fig. 7), Clusters 2, 3 and 4 from the SHGF and the Hanlé
geothermal well (BFGUR, 1999) are partially equilibrated and more
and more shifted towards full equilibrium, suggesting that chemical
geothermometers can be used with care to estimate geothermal reservoir temperature (Fournier and Truesdell, 1973; Fournier and
Potter, 1979; Giggenbach, 1988). On the other hand, samples from
Cluster 1 and boreholes from the SHGF and the Hanlé geothermal
well, sampled during this study are more shifted towards the magnesium corner and thus has not equilibrated with local volcanic rocks.
In the square plot (Fig. 8), the same trend for SHGF samples is
shown as a vertical path which departs from an almost horizontal
line that joins a non-equilibrium curve for isochemical dissolution
of local basalts and a full equilibrium curve at 35 °C (the mean local
air temperature). For example, warm and cold water from the
Obock borehole (not shown, Awaleh et al., 2015a) and samples
from the Hanlé geothermal well sampled in this study plot onto the
horizontal line, suggesting that they are not useful for cation
geothermometry. On the contrary, and according to Giggenbach
(1988), the vertical path should suggest that mixing occurred
between shallow and deeper aquifer and that the Na-K
Galiceela (#90) hot spring (Fouillac et al., 1989), but a chloride versus
water isotope plot reveals a possible dilution trend, as shown by other
thermal springs in the Asal geothermal field (Fig. 6). Moreover, in
these diagrams it is evident how Sakalol springs are generated by a
mixing between three end-members: meteoric water with low chloride
concentration (cluster 1), meteoric origin fluids from the surrounding
grabens (T.A.D.D. field and Hanlé springs with approximatively
δ18O = b− 1‰ and δ2H b − 10‰) and rift water with high chloride
concentration (Asal geothermal wells). The graben and rift components
appear to be interrelated because both could be influenced by alluviallacustrine evaporates resulting in convergent chloride shifts. For
examples, the Dobi graben is characterized by the presence of salt pan,
remnants of a larger lake that filled it about 2000 ya (Kebede et al.,
2008; Williams, 2016). However, supposing that chloride in cluster C4
(Galiceela sample; Fig. 6) derive mainly from a deep fluid, a contribution
of approximatively 7% from Asal geothermal water can be inferred; the
remaining 93% is attributable mainly to a regional fluids, whereas a
fresh meteoric input should be dominant in the other clusters.
4.4. Geothermometry
The Na/1000–K/100–Mg0.5 ternary plot of Giggenbach (1988) can
be used to discriminate mature waters, which have attained equilibrium with relevant hydrothermal minerals, from immature waters
and waters affected by mixing and/or re-equilibration along their
circulation path, providing indications of the suitability of the waters
Na/1000
Abhe
160
full equilibrium
140
100
200
#69
Allallobeda
Asal
Hanlé
#58
300
Tendaho
Dobi
partial equilibrium
160
140
200
immature waters
AC
#35
AB
K/100
Mg
well-borehole
Cluster 1
Sakalol
Cluster 2
BFGUR 1999
Hanlé
geothermal
wells
this study
Cluster 3
Cluster 4
Houssein et al. 2010
Dobi
waters with Cl > 1 g/l
seawater
Fig. 7. Trilinear geothermal diagram (samples data in mg/l; Giggenbach, 1988). AC and AB: isochemical dissolution of average crust and basalt, respectively (Giggenbach, 1988).
Temperatures of full equilibrium and immature waters/partial equilibrium curve are in degrees Celsius (°C). The Dobi graben springs for Cl N1 g/l have been specified (triangles).
Samples data as in Fig. 3.
38
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
seawater
1.0
Lake Asal
20
40
0.9
tMg-Ca
Isochemical
Dissolution
Basaltic
Trap
Obock
0.8
Basaltic
Lava
#35
Acid Tuff
60
0.7
Agna
Full
0.5
Rhyolite
um
Equilibri
10Mg/(10Mg+Ca)
#36
0.6
80
Minkileh
0.4
Dobi
100
0.3
0.2
tMg-Ca
120
tK-Na
Oued-Kalou
Tendaho
140
320
160
0.1
0.0
0.0
Danab
200
180
Abhe Allalobeda
0.1
0.2
0.3
260
220
0.4
340
280
240
0.5
0.6
300
Asal
0.7
0.8
0.9
1.0
10K/(10K+Na)
waters
rocks
literature
this study
borehole
Asal
Cluster 1
Sakalol
Cluster 2
Cluster 3
Hanlé
geothermal well
hot spring
Dahla Basalts
(Gasse 1987)
geothermal well
Asal Basalts
(Pinzuti et al. 2003)
borehole
hot spring
Cluster 4
Dobi
Various Type
(Bosch et al. 1977)
water with Cl > 1 g/l
Fig. 8. Square geothermal diagram (samples data in mg/l; Giggenbach, 1988). Water samples data as in Fig. 3.
geothermometer (tNa-K on Fig. 8) could give more coherent results in
comparison with geothermometers that use Ca and Mg ions (tCa-Mg
on Fig. 8). Indeed, minor horizontal and vertical shifting along the
vertical path is probably due to water re-equilibration at more
supergenic conditions. This is confirmed plotting the SHGF samples
in activity diagrams which testify equilibrium with zeolites
(Fig. 9a), clays (smectites as saponites and montmorillonites) and/
or carbonates (Fig. 9b). The SHGF samples pertaining to the different
clusters are more and more shifted along the vertical path and towards the full equilibrium curve, on which Cluster 4 samples
(Galiceela spring) fall quite close to the 160 °C temperature point
(Figs. 7 and 8). However, it should be noted that the equilibrium
curve represents a re-crystallization of a mean crustal rock from outdated thermodynamic data (Giggenbach, 1988; Cortecci et al., 2005).
Moreover, the position over to the equilibrium curve on the square
plot in comparison with the ternary geothermometric diagram
representing the Djibouti (Abhe, Hanlé, Sakalol-Harralol and Asal)
and Dobi graben (Cl N 1 g/l) hot waters could testify to the addition
of calcium ion from secondary minerals, for example by interaction
with sulfates (Fig. 8). Therefore, in the following paragraph, classical
and new chemical and isotope geothermometers were used and
compared in order to obtain the best estimation of the reservoir temperature of the SHGF.
4.4.1. Chemical geothermometers
The Na-K geothermometer is related to the variation of sodium and
potassium in thermal waters due to ion exchange of these elements between coexisting alkali feldspars (e.g. Nicholson, 1993). The Na–K
geothermometer of Giggenbach (1988), based on thermodynamic
data for albite and K-feldspar, gives an equilibrium temperature of
deep parent fluids of about 132–184 °C (Supporting information file).
Furthermore, the Na–K geothermometers of Fournier (1979) and
Arnórsson et al. (1983) estimated the temperature of the geothermal
reservoir respectively at about 117–173 °C and 117–168 °C (Supporting
information file). These temperatures agree with those obtained with
Na–K–Ca geothermometer: 130–168 °C (Supporting information file).
Despite this agreement, it should be noted that most of the alkali
geothermometers could not give precise results because they are
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
e
lin ite
s
oc
icr illip
e
te
h
di
lin
-p
a
oc lan
N
icr heu
e
m
t
i
Na ite
ps
e illi
ps
sit ph
ite
illi
h
llip Nand
i
h
-p
la
p
K
eu
Kh
Na
6.0
5.5
25°C
without Na-heulandite
with Na-heulandite
80°C
without Na-heulandite
with Na-heulandite
m
+
+
log[K /H ]
5.0
4.5
K-beidellite
#90
4.0
Na-beidellite
3.5
A
3.0
5.0
5.5
6.0
6.5
+
7.0
7.5
+
log[Na /H ]
14.0
13.5
25°C
carbonates
silicates
80°C
carbonates
silicates
e
sit s
ne -di
ag te
m mi e
lo it
do calc
13.0
Mg-saponite
2+
+ 2
log[Mg /(H ) ]
12.5
Ca-saponite
12.0
Mg-montmor
11.5
e
sit
ne
ag
m
11.0
is
-d
ite
om
l
do
te
Ca-montmor
lci
ca
scolecite
10.5
#90
kaolinite
10.0
10.5
11.0
11.5
39
basalt-seawater and Cl N 0.3 M equations give mean values lower than
the major cations geothermometers: 112 ± 71 °C (median 89 °C) and
107 ± 49 °C (median 92 °C), respectively. The negative results obtained
by the first equation and the higher temperature obtained by other two
could be due to the combined effects of mixing and re-equilibration
with clay minerals (as testified by activity diagrams, Fig. 9). In particular,
clay adsorption could be particularly effective for the lithium concentration, Li being a trace element. A temperature of 174 °C was obtained for
the Galiceela springs by the Cl N 0.3 M equation. Considering the relatively high lithium and chloride concentrations of this water and the uncertainty of ±20 °C related to these geothermometers (Sanjuan et al.,
2014), this results should be considered more reliable than that obtained for other springs.
Among several existing silica geothermometers, those proposed by
Fournier (1977) for quartz, and that of Arnórsson et al. (1983) for chalcedony, have been used successfully both for alkaline thermal waters
and waters with T b 180 °C (Michard and Beaucaire, 1993; Marques
et al., 2003; Asta et al., 2012; Fournier, 1991). This agrees with the supersaturation state of the studied waters with respect to those minerals:
0.56 ± 0.13 for chalcedony and 0.82 ± 0.13 for quartz (see supporting
information). Therefore, only chalcedony and quartz have been chosen
as silica geothermometers. In the present study, the quartz
geothermometers of Verma (2001) estimate the temperature of the
SHGF geothermal reservoir in the range of 111–148 °C, about 3° higher
than Fournier's equations and mean temperature between 133–136 °C
(Supporting information file). On the other hand, the reservoir temperature values estimated by the chalcedony geothermometers are in the
range 94–127 °C (Supporting information file). It should be noted that
most of the waters sampled are grouped between the chalcedony and
amorphous silica equilibrium lines (Fig. 10). In particular, some borehole waters at the SHGF showed high dissolved silica concentrations
(~ 100 mg/l) and plot near to the amorphous silica saturation at low
temperature. However, globally the waters studied are undersatured
with respect to amorphous silica (SI = −0.29 ± 0.09; see supporting
information) and their outlined path suggests mixing with the local
and shallow groundwater after a conductive cooling (Fig. 10). Also for
silica, the highest temperatures of 148–149 °C were obtained for the
Galiceela sample, which suffered less cooling and mixing. For the
same spring, temperatures of 152 °C and 158 °C were obtained by two
equations calculated from the results of a leaching experiment performed on local basalts (trap and lava; Bosch et al., 1977):
B
12.0
12.5
2+
13.0
13.5
14.0
+ 2
T ð °CÞbasaltic trap ¼ ½−1576:085=ðlogC–5:799Þ–273
ð1Þ
T ð °CÞbasaltic lava ¼ ½−2070:977=ðlogC–6:894Þ–273
ð2Þ
log[Ca /(H ) ]
Fig. 9. Na-K (A) and Ca-Mg (B) activity diagrams. The Geochemist's Workbench (Bethke
and Yeakel, 2008) along with Thermoddem (release 2014; Blanc et al., 2012) and thermo.
com.v8.dat (B) thermodynamic datasets have been used to draw the solid phase limits in
diagrams (A) and (B), respectively. Activity of the dissolved species was calculated by
PHREEQCI (Parkhurst and Appelo, 2013).
calibrated more than 30 years ago, that is using thermodynamic data to
be updated (Cortecci et al., 2005). Therefore, as a first approach, it would
be better to choose more recently calibrated equations or pure thermodynamic data (i.e. mineral solubility product and activity of dissolved
species) by using updated datasets. In fact, for the Galiceela sample
(#90), the value of 146 °C obtained using the Na-K equation of Verma
and Santoyo (1997) is 12 °C lower than Giggenbach’s one (160 °C, as
also is estimable from Fig. 8) but more similar to that of 145 °C obtained
by Na-K-Ca geothermometer (Supporting information file).
Several Na-Li geothermometers have recently been proposed and/or
recalibrated (Sanjuan et al., 2014). However, what should be the most
appropriate for our samples (Cl b 0.3 M, equation 7 in Sanjuan et al.,
2014) is unproductive because it gives negative results, whereas
where C = SiO2 in mg/l.
Temperatures by a pure thermodynamic approach were estimated
using PREEQCI and the full composition of the Galiceela samples. The activity of the dissolved species of this latter was recalculated at specific
temperature up to quartz saturation. This was obtained at 151 ± 1 °C,
which is consistent with the result obtained from the “basaltic trap”
equation. At that temperature, the new activity of sodium and potassium was used to calculate the temperature of equilibrium between the
alkali feldspars:
albitelow þ Kþ ¼ K−feldspar þ Naþ
The logK of this reaction was calculated by using Rxn tool of The
Geochemist’s Workbench® 7.0.6 software, version 7.0.6, and thermo.dat
logK ¼ 3:462−0:0194 T þ 6:449e−5 T2 −1:214e−7 T3
þ 9:581e−11 T4
ð3Þ
40
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
mean local air
temperature (35 °C)
A2
hou
Dobi
am
orp
120
C1
Abhe-1
ss
ilic
a
140
conductive
cooling
Q3
Minkileh
A1
100
Abhe-2
80
B1
60
y
on
Q1
z
art
ba
40
sa
ed
alc
ch
B2
lts
SiO2 (mg/l)
A3
qu
Q2
20
C2
0
10
20
30
40
50
60
70
80
90
100
110
120
130
140
150
160
Temperature (°C)
well-borehole
Cluster 1
Sakalol
Cluster 2
Asal
hot spring
Hanlé
geothermal well
springs
Cluster 3
Cluster 4
Fig. 10. Dissolved silica vs. temperature. Sakalol data (diamonds) from this study. Asal (circles) and Hanlé (squares) from the literature (see Fig. 3 for references). Squares with sharp edges
are Hanlé data from this study. Curves represent the solubility of the different polymorphs: amorphous silica A1 (Verma, 2001), A2 (Fournier, 1977), A3 (Gunnarsson and Arnórsson,
2000); chalcedony C1 (Fournier, 1977), C2 (Arnórsson et al., 1983); quartz Q1 (Verma, 2001), Q2 (no steam loss, Fournier, 1977), Q3 (steam loss, Fournier, 1977). Equation describing
the silica obtained from dissolution at different temperature of local trap and lava basalts (B1 and B2, respectively) have been calculated from the data of Bosch et al. (1977). Means
and standard deviations for waters from Abhe 1 and 2 (small and great hydrothermal chimneys, respectively; Awaleh et al., 2015b) and the Dobi graben (Ali, 2005; UNDP, United
National Development Program, 1973) are also shown for comparison.
or thermo.com.v8.r6+.dat database
logK ¼ 3:489−0:01952 T þ 6:602e−5 T2 −1:303e−7 T3
þ 1:138e−10 T4
4.5. Dissolved sulfate: origin traced by isotope
ð4Þ
In the equations above, log K = log[Na+] – log[K+], where square
parentheses are the activity of the specie obtained from PHREEQCI,
and the temperature in degree Celsius.
Eq. (3) gives a temperature of 134–141 °C, whereas the equation
Eq. (4) 137–144 °C. Therefore, a temperature range of 143 ± 13 °C
would be more realistic for the deep SHGF reservoir.
It is interesting to note the compositional transition of the Dobi graben springs with Cl N 1 g/l: were similar to the Allalobeda waters during
first analysis (UNDP, United National Development Program, 1973)
and then (Ali, 2005, 2011; Panichi, 1995) progressively to AbheSakalol-Hanlé reservoir (Fig. 8). Such a change was probably related to
an earthquake sequences (Noir et al., 1997) which enhanced permeability and the communication between fluids at the Ethiopia-Djibouti
border.
The SO4/Cl vs SO4 plot is very useful to trace sulfate origin and its fate
in groundwater (e.g. Kehew, 2001). In Fig. 11 C1 to C2 samples increase
concentration with decreasing SO4/Cl ratios, which is the opposite behavior of solid sulfate dissolution. Moreover, and similarly to that of
the anion ternary plot, the position of some C1 and C2 springs in the diagram depicts the evolution path towards the Galiceela spring (C4
group; Fig. 11). However, the increase of sulfate concentration is in accordance with the evolution of the mean gypsum/anhydrite saturation
indices of the SHGF cluster groups (an half of C2 samples and almost
all C3 samples fall near sulfate dissolution at 143 °C and constant NaCl concentration; Fig. 11), whereas the decreasing SO4/Cl ratio testifies
to a deep chloride input. In particular, and in correspondence with the
anhydrite saturation point at the mean deep temperature of 143 °C, almost vertical trends (constant sulfate concentration and variable SO4/Cl
ratios) are revealed also in hot springs from neighboring systems. For
example, springs from the Hanlé plain show increasing SO4/Cl ratios
departing from Allallobeda-Abhe-Dobi mean values. This result is in
agreement with the suggestion of a meteoric input from shallow sedimentary formations (SO4/Cl up to 3 meq/l; data from Houssein
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
41
0.6
bicarbonate waters
interacting with
sedimentary formations
0.5
Minkileh
gypsum
saturation
0.3
2-
-
SO4 / Cl (meq/l)
0.4
Allallob.
sulfate dissolution
(143 °C)
Agna
Dobi
0.2
Abhe
meteoric
input
Danab
0.1
sulfate dissolution
(25 °C)
anhydrite
saturation
Tendaho
seawater
Obock
#75
marine(?)
input
0
10
100
1000
10000
2-
SO4 (mg/l)
this study
literature
well-borehole
Asal
Sakalol
geothermal well
geothermal well
Cluster 1
Cluster 2
Oued-Kalou
hot springs
Hanlé
borehole
hot-warm-cold springs
Cluster 3
Cluster 4
Dobi
Cl > 1 g/l waters
Fig. 11. Sulfate/chloride ratio vs. sulfate concentration. Dotted curves represent calcium sulfate dissolution at different chloride concentrations (average of C1 and C2 clusters) and
temperatures, that is Cl = 15 mmol/kg H2O – T = 25 °C and Cl = 62 mmol/kg H2O – T = 143 °C, respectively. Bricked field represent bicarbonate water interacting with sedimentary
formation in the Hanlé graben (see Fig. 4). Mean ± standard deviation of waters from other geothermal fields in the Afar is also shown for comparison (see the Fig. 3 caption and
main text for references).
(2010). In a similar way: i) deep and shallow borehole waters from
Hanlé are grouped on the low temperature gypsum dissolution curve;
ii) Oued-Kalou springs waters at Asal show high SO4/Cl ratios and
group near to Dobi and Abhe mean values when meteoric input is
high, whereas the ratio is lowered and merge to the values of the Asal
geothermal wells when a marine input prevails. A similar deep input
of marine origin, i.e. a “recycled” seawater in the Asal rift geothermal
system, could be deduced for Galiceela samples, which fall near to the
back-extrapolation curve of the Obock samples. Alternatively, solid sulfate dissolution is enhanced by a high chloride content due to interaction with evaporitic minerals, as is taking place for the Dobi graben
samples with a relatively high chloride concentration (Fig. 11). In
support to the second hypothesis, it should be noted that the sulfur isotope composition for dissolved sulfate at Galiceela (+11.4‰) is not different from the mean value for the SHGF springs waters (+ 11.4 ±
0.6‰) and very different from marine sulfate (+21‰). In Fig. 12a, it is
evident that most of the sampled waters are grouped near the 34Senriched component represented by the solid and dissolved sulfate
from the Asal geothermal area (11.4‰ b δ34S b 16.6‰; Fouillac et al.,
1989; Gasse and Fontes, 1989). A closer look at the samples' distribution
in Fig. 12a reveals also that: i) 34S enriched component has a more restricted range of 11‰ b δ34S b 14‰, this latter is comparable with sulfate
from Lake Abhe hot springs waters (Awaleh et al., 2015b); ii) fresh waters of Cluster 1 are more shifted towards a line joining Abhe springs
42
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
Lake Asal
18
A
17
34
2-
δ S(SO4 ) (‰ vs. V-CDT)
16
15
14
13
12
11
10
Oued-Kalou #75
9
#35
8
0
2
4
6
8
10
12
14
16
18
20
3
2-
(1/SO4 ) x 10 (mg/l)
literature
this work
Abhe
(hot springs)
well-borehole
Cluster 1
Cluster 2
Cluster 3
Cluster 4
Asal Rift
(solid and dissolved sulfates)
Asal hot springs
18
30
0°C
250°C
B
17
.5
16
=0
15
2
350°
C
H2 S/S
11
HS
2 /S
O
O2 = 1
x10 -4
12
C
18
13
200°
δ O(SO2-4 ) (‰ vs. V-SMOW)
14
10
9
40
#75
0°C
8
O
H 2S/S 2
7
=1
Lake Asal
#35
6
5
4
3
2
1
sulfide oxidation
0
0
2
4
6
8
34
10
12
14
16
18
2-
δ S(SO4 ) (‰ vs. V-CDT)
Fig. 12. Sulfur isotope vs. sulfate concentration (A) and oxygen vs. sulfur isotope composition of dissolved sulfate (B). In both diagrams, solid and dissolved sulfate data from Asal Rift
basalts (gray field) and Lake Asal are from Fontes et al. (1979), Fouillac et al. (1989) and Gasse and Fontes (1989). Dashed lines depict possible mixing or evolutive trends. In (B), the
grid represents sulfate composition after disproportionation of local magmatic SO2 (see text for details).
waters and a 34S-depleted component (δ34S b 11‰), represented by the
Dora borehole water (#35); iii) value of δ34S N 16.6 are typically found
where a seawater intrusion prevails (e.g. Manda spring or Lake Asal,
both in the Asal Rift area; Fontes et al., 1979; Gasse and Fontes, 1989).
Gypsum mounds analyzed in the Asal Rift showed values of 1.08‰ b
S b 16.3‰, which were interpreted as due to disproportionation of
magmatic SO2 but not verified by calculation (Moussa et al., 2016).
Fig. 12b helps verify this hypothesis and to better understand the sulfate
34
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
source in the aquifer in the SHGF. The Asal sulfate field intercepts a grid
that represents the isotope effect due to SO2 disproportionation at
200–350 °C. The grid was calculated using a similar approach to that of
Rye et al. (1992), but using the disproportionation model and
0
103lnα(SO2−
4 -S ) equation of Kusakabe et al. (2000) and mean values
for 103lnα(SO24 −-H2S) obtained from Sakai (1968), Robinson (1973)
and Ohmoto and Lasaga (1982). A primary magmatic value of
δ34S(ΣS) = −2.78‰ (Allard, 1979) was maintained for all calculations,
whereas the H2S/SO2 ratio was varied from 1 × 10−4, typical of the
Asal magmatic gas (Allard, 1979) up to 1 in accordance with the SO2
switch to H2S and SO24 − during disproportionation and/or H2S/SO2
ratio increase at near surface conditions (e.g. Kusakabe et al., 2000;
Wallace and Edmonds, 2011). At the same time, to simulate this transition during the modeling calculation, the water isotope composition was
gradually changed from δ18O(H2O) = +6.0‰ (according to the basaltic
composition of the magmatic source; Fig. 5) to δ18O(H2O) = +0.33‰
(mean value of the Asal geothermal water; Correia et al., 1985;
D'Amore et al., 1998). The comparison between the grid obtained, Asal
field and SHGF samples of this study reveals some common features
(Fig. 12b). The coupled low 34S - high 18O values of the Asal sulfate
field agree with a SO2-disproportionation at 250–300 °C, according to
the temperature of the deep reservoir (D'Amore et al., 1998). The high
temperature low 34S - high 18O couple of the Asal field agree with the
values measured in the deep basalt and with the δ34S of the hot springs
from the Asal geothermal area (Bosch et al., 1977; Fouillac et al., 1989).
40
18
#90
16
Sakalol reservoir
(143 °C)
14
Abhe
hot
Lake
Asal
springs
12
-H
O
2
4
-
-H
4
HS
O
SO
Ca
3
10 lnα
8
2
O
#75
10
2-
SO
6
4
-H
O
2
#35
M.E.R.
4
2
Allalobeda
(79÷95 °C)
0
-2
3
4
5
6
7
6
8
9
10
11
12
2
10 /T (kelvin)
this study
Sakalol
well-borehole
Cluster 1
Cluster 2
Cluster 3
Cluster 4
The low temperature high 34S – low 18O couple agrees with the value
measured in the surface basalts and gypsiferous deposits at Asal
(Fouillac et al., 1989; Gasse and Fontes, 1989). All Sakalol water samples
of this study fall within the grid and most of them near the Asal field.
Therefore, we suppose that water samples at the SHGF acquired a
disproportionated magmatic sulfate similar to those of the Asal area
interacting with sulfate minerals which may be included in the basaltic
rocks, fumarolic condensates or recycled in sedimentary phases. Supposedly dissolved or deposited sulfates, whose values fall out of the grid,
have an origin different from SO2-disproportionation of magmatic gas.
For example, in Fig. 12 this is the case for Asal sulfates with δ34S N 17‰
and δ18O b 9‰ which were influenced by Holocene seawater (e.g. Lake
Asal; Fontes et al., 1979; Gasse and Fontes, 1989). On the other hand,
and according to the sulfur general model at SHGF, the Cluster 1 and
the Dora borehole waters were more subjected to the oxidation of sulphide driven by atmospheric oxygen, this latter coming from fresh basalt
and/or endogenous gas (i.e., in a similar way to that of steam-heated waters; Boschetti et al., 2003; Bayon and Ferrer, 2005).
Finally, in Fig. 13, the temperature and SO4-H2O isotope fractionation factors of the SHGF springs waters and deep source are compared
to theoretical equilibrium fractionation and the literature data for Abhe
(Awaleh et al., 2015b) and Hanlé (Aquater, 1986) springs. Samples distributions confirm common water and sulfate sources. Despite similarity of the temperatures in the Hanlé-SHFG system and the M.E.R.
(Craig et al., 1977), these latter show lower oxygen fractionation values
mainly due to higher δ18O(H2O), as shown in the water isotope diagram
(Fig. 5). Unfortunately, sulfate isotopes data for waters and minerals
from the Dora graben are not available.
4.6. Halogens (F, Cl, Br), boron and 11B/10B ratio
80
12
16
0
0
0
20
0
24
28
0
Temperature (°C)
43
literature
Hanlé springs
(Aquater 1986)
Fig. 13. Sulfate-water oxygen isotope fractionation as 103lnα (‰) vs. temperature (Kelvin
and Celsius) for Sakalol springs (diamonds; this study). Full equilibrium lines are shown
(see Boschetti, 2013 for equations). Like Abhe deep geothermal water (Awaleh et al.,
2015b), the temperature obtained by chemical geothermometers for Galiceela hot
spring (#90) is compatible with an oxygen isotope equilibration between anhydrite and
water. Data from M.E.R. hot springs, with focus on Allalobeda (mean ± standard
deviation, UNDP, United National Development Program, 1973); Abhe (Awaleh et al.,
2015b) and Hanlé (Aquater, 1986; squares) spring waters are also shown for
comparison. The vertical dashed line represents mean local air temperature (T = 35 °C),
but arrows possible mixing or evolutive trends.
Like in other waters of the so called “Fluoride Belt” of the East African
Rift (Gupta and Ayoob, 2016), the mean fluoride concentration of
4.13 ± 1.70 mg/l in the spring waters of the SHGF is higher than the
1.5 mg/l concentration limit for drinking water (WHO, World Health
Organization, 2011). The SHGF mean value is higher than that for the
Tendaho-Allallobeda waters, similar to that for the Dobi graben and
Asal waters, lower than the concentration shown for M.E.R. hot springs
(28.4 ± 25.6 mg/l; Rango et al., 2013) and much lower than those for
the geothermal waters in that area (e.g. 71 ± 17 mg/l in the AlutoLangano Geothermal Field; Teklemariam and Beyene, 2001) which are
Na-bicarbonate waters (Fig. 14a). The cause of the high concentration
in M.E.R. waters was related to low calcium concentration in solutions
and interaction with fluorite, which could be present in sedimentary
and volcanic formation and is more soluble than other fluoridebearing minerals (Gizaw, 1996; García and Borgnino, 2015). Calculated
saturation indices show fluorite undersaturation for SHGF springs with
a very high standard deviation: −1.32 ± 1.13. The reason for this high
scatter is probably also the cause of the different concentrations from
those of the M.E.R. hot waters, that is the simultaneous dissolution of
calcium sulfate minerals which cause fluorite precipitation due to a
common-ion effect (e.g. Appelo and Postma, 2007). Scattering is higher
in the diluted waters of the SHGF C1 cluster and is as also shown by the
Hanlé geothermal wells. Fluorine concentrations are due to interaction
with sediments or volcanic ash or scoria of the shallow and recent formations (e.g. Ghiglieri et al., 2012).
Unlike fluoride, boron isotopes are good geochemical tracers because of the high mobility of this element during water-rock interaction
processes (Hoefs, 2015). Boron is preferentially hosted in
phyllosilicates, whereas the B concentrations of common mantle and
crustal minerals (except tourmaline) are low. The low boron concentrations of the Tadjourah Gulf basalts (0.2–0.8 ppm and −4 ÷ −15‰ δ11B;
Chaussidon and Marty, 1995) confirm this theory. B versus Cl and δ11B
versus Cl/B are useful diagrams for the study of geothermal systems
(Aggarwal et al., 2000; Bernard-Romero et al., 2010; YuanYuan et al.,
2014). To the best of our knowledge there are no published δ11B data
44
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
Cl (mg/l)
10
10
100
1000
100000
M.E.R. hot springs: (F = 28.5 ± 25.6 mg/l)
(Rango et al. 2013)
M.E.R. groundwaters
(Rango et al. 2013)
9
10000
8
7
F (mg/l)
6
5
4
Dobi
3
2
1
WHO (2011)
Allallobeda
seawater
Abhe
F limit in groundwater
Obock
Tendaho
A
#35
M.E.R. geothermal waters
(Aluto-Langano)
0
10
9
8
7
B (mg/l)
6
M.E.R. hot springs
(Rango et al. 2013)
seawater
5
Tendaho
Obock
4
3
Abhe
Allallobeda
basalts
2
Asal
(fum. condensate)
1
Dobi
#35
B
0
10
M.E.R. groundwaters
(Rango et al. 2013)
100
1000
10000
100000
Cl (mg/l)
this study
well-borehole
literature
Asal
geothermal well
hot spring
Cluster 1
Sakalol
geothermal well
Cluster 2
Hanlé
Cluster 3
borehole
hot spring
Cluster 4
Dobi
Cl > 1 g/l waters
Fig. 14. Chloride concentration vs fluoride (A) and boron (B) concentrations. Literature data sources as in Fig. 3. In (B) “basalts” curves represent typical B/Cl ratio in basaltic rocks
(Arnórsson and Andresdottir, 1995; Reyes and Vickridge, 1996), which are similar in Aluto-Langano geothermal waters from the M.E.R. (Gianelli and Teklemariam, 1993; Teklemariam
and Beyene, 2001). Dashed, dotted and dash-dot depict Asal-Tendaho, seawater-meteoric and Asal geothermal water Asal fumarole (Virkir-Orkint, 1990) binary mixing lines. In both,
the arrow depicts a possible evolutive trend for Sakalol hot spring waters.
for geothermal waters in the African Rift at the time of writing. In
Fig. 14b, it is evident that the boron concentration increases from Cluster 1 to Cluster 4, like in other Afar rift hot waters. In particular, the
boron concentration of most Cluster 3 and Cluster 4 (Galiceela) waters
were relatively high compared to Cluster 1 waters and more shifted towards seawater-derived fluids such as Asal or Obock. This is also
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
45
A
marine source
Lake Asal
Asal geothermal reservoir
11
δ B (‰ vs SRM951)
35
25
15
d
5
lan
ice
non-marine
source
-5
45
basalts
-15
1
10
100
1000
10000
Cl/B (mg/l)
this study
literature
well-borehole
Asal
Sakalol
Cluster 2
geothermal well
Cluster 3
Hanlé
Cluster 4
0.1
B
geothermal well
hot spring
Cluster 1
borehole
hot spring
hydrothermal
#35
basalts
Allallobeda
Abhe
0.01
Danab
B/Cl (mol)
sewage
Dobi
seawater
evaporation
0.001
agriculture drainage
evaporites
0.0001
0.0001
#74
Lake
Asal
0.001
seawater
intrusion
0.01
Br/Cl (mol)
Fig. 15. Br/Cl vs B/Cl (A) and Cl/B vs δ11B (B). In (A), fields represent different geothermal waters worldwide from Aggarwal et al. (2000), Bernard-Romero et al. (2010) and YuanYuan et al.
(2014). Exception is provided by the Lake Asal and Asal geothermal reservoir fields, whose Cl/B ratios only are shown (Bosch et al., 1977; D'Amore et al., 1998; Sanjuan et al., 1990) because
boron isotope ratios are not available. The basalt field data was adapted to local rocks for isotope ratio (Chaussidon and Marty, 1995) and Cl/B ratio from literature data (see text and B vs. Cl
plot). In (B), fields and paths are from Vengosh (2014), except “basalt” (Arnórsson and Andresdottir, 1995; Möller et al., 2016). See Fig. 3 for the references on the sample data from Djibouti
(Abhe, Asal, Hanlé) and the Ethiopian Afar (Allalobeda, Danab, Dobi). In (B), fields represent different geothermal waters worldwide (Aggarwal et al., 2000; Bernard-Romero et al., 2010;
YuanYuan et al., 2014).
confirmed by higher ratio Cl/B ratios of the SHGF (median = 853 mg/l)
than those of the hot springs and geothermal wells from M.E.R., these
latter ratios are more similar to the low Cl/B typical of basalts (30–120
by weight, Fig. 14b; Arnórsson and Andresdottir, 1995; Reyes and
Vickridge, 1996). As shown in Fig. 15, this could be interpreted in several ways such as a seawater origin for these elements or an involvement
of a fluid similar to the Asal geothermal fluid in the SHGF geothermal
circuit. Taking into account the involvement of seawater in the Asal geothermal system (D'Amore et al., 1998), the assumption of one hypothesis does not exclude the other. In fact, in spite of the different boron
concentrations in clusters, all samples are shifted towards a marine or
a geothermal-marine source. The relatively high 11B value for the
46
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
4.7. Strontium and carbon isotopes
SHGF waters could be due: i) B delivered by gas, but this should be a
very minor contribution, considering the low concentration of boron
in Asal and Hanlé fumaroles (Aquater, 1986; Virkir-Orkint, 1990); ii)
boron fractionation due to preferential inclusion of 10B in carbonate
and clay minerals (e.g. Hemming et al., 1995). This hypothesis is more
probable considering the equilibrium or slight supersaturation with respect to aragonite and calcite, respectively, and the clay-zeolites activity
diagrams (Fig. 9). The probable secondary fractionation effects and the
lack of δ11B data for the surroundings geothermal reservoirs (e.g. Asal,
Abhe, Hanlé, Dobi) which may hide some informations, the use of a
more discriminating diagram such as Br/Cl vs B/Cl on a molar basis
(Shikazono, 2003; Vengosh, 2014) could be useful (Fig. 15). Most of
the samples from the SHGF are grouped within or in the low right corner of the hydrothermal field suggested by Vengosh (2014). Abhe,
Dobi and Allallobeda geothermal waters depart from this field, but Cluster 1 follows the grouping depicted by the mean and standard deviation
values, suggesting that the right limit for a hydrothermal field should be
extended slightly. This is also confirmed by the results for the Dora sample, which falls within the ratios values for basalts (Br/Cl = 0.0013 to
0.0023 in Möller et al., 2016; B/Cl = 0.022 to 0.1 of alkali basalts in
Arnórsson and Andresdottir, 1995). Finally, the rough linear trend
shown by Clusters 2, 3 and 4 confirm the involvement in the SHGF of
a geothermal water of marine origin, similar to the Asal geothermal
water (Fig. 15).
The 87Sr/86Sr ratio determined in water reflects the distinct isotope
composition of the interacting rocks (Goldstein and Jacobsen, 1987;
Drever, 1997). Few water 87Sr/86Sr data are available for Djibouti
(Sanjuan et al., 1990; Dekov et al., 2014; Awaleh et al., 2015b; Awaleh
et al., 2016). The ratios of the investigated water samples range from
0.70365 to 0.70626 (Table 2). This data is compared with the 87Sr/86Sr
values of local lithotypes and waters available in the literature in order
to define the nature of the interacting rocks and possible relationships
(Fig. 16). Sr isotope values for basalt samples from the Republic of
Djibouti show a considerable variation from 0.70309 to 0.70664,
where the higher values correspond to older rocks (10 to 25 My) that
included a subcontinental lithospheric component (Barrat et al., 1993;
Deniel et al., 1994; Faure, 2001; Vidal et al., 1991).
The relatively low 87Sr/86Sr ratio (0.70365; Table 2) observed in C4
thermal waters indicates that Sr is predominantly leached from the underlying volcanic rocks, most likely from the Stratoid Series that dominates the bedrock in the Southwestern region of the Republic of
Djibouti (Vidal et al., 1991; Barrat et al., 1993). According to a shift towards the Sr isotope ratios of Quaternary Stratoid basalts (Fig. 16), the
87
Sr/86Sr value for C4 thermal water from the SHGF is slightly lower
than the Sr isotope value for Djibouti geothermal waters ranging from
0.703809 to 0.70463 (Awaleh et al., 2015b; Sanjuan et al., 1990). In a
87
0.703
100
0.704
«m
10
eo
et
»
ric
0.705
in
orig
86
Sr/ Sr
0.707
0.706
gs
sprin
hot
0.709
0.710
#35
Abhe
1
1/Sr (mg/l)
0.708
Oued-Kalou
K2
alluvial/uncofined
groundwaters
(Awaleh et al. 2016)
Oued-Kalou
K1
seawater
0.1
Korili
«m
0.01
Afar
Plume
(0.70350)
0.001
e»
a rin
ho t
origin
Red-Sea brines
(Pierret et al. 2001)
springs
Lake Asal
Dahla
basalts
(0.70411)
Quaternary
Basalts
(0.70366)
low-T travertine
(Dekov et al. 2014)
0.0001
this study
literature
well-borehole
Cluster 1
Sakalol
Cluster 2
Asal
geothermal well
hot spring
Cluster 3
Cluster 4
Fig. 16. 87Sr/86Sr vs. 1/Sr. Means and standard deviations for Djibouti basalts (Faure, 2001) and Abhe hot springs (Awaleh et al., 2015b) are shown for comparison. Curves represent binary
mixing of the hot waters in Sakalol (this study) and Asal (Sanjuan et al., 1990) which represents geothermal waters of meteoric and seawater origin. It is interesting to note that the OuedKalou spring waters from Asal plot in the middle (dotted line is the best fit for these springs), whereas Red sea brines (Pierret et al., 2001) plot between seawater and Lake Asal water
(Sanjuan et al., 1990). The dashed line encloses non-thermal groundwater samples from alluvial and Dahla basalt aquifers (Awaleh et al., 2016).
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
47
-
H CO3 (mg/l)
seawater
100
Allallobeda
(Craig et al. 1977)
200
300
400
500
600
700
800
0
-2
Agna
African-Rift
high-T springs
(Darling et al. 1996)
mantle
CO2
M.E.R. hot springs
(Bretzler et al. 2011)
-4
-6
-8
Dora (#35)
Water-atmospheric CO2
equilibrium at 35°C
-10
water-silicate interaction
(open system)
13
-12
#36
-14
rift-escarpment CO2
(Bretzler et al. 2011)
δ C(DIC) (‰ vs V-PDB)
0
Asal geothermal reservoir
(D’Amore et al. 1988)
-16
low-T groundwater
A
-18
Minkileh
120
seawater
100
Abhe hot spring
(Fontes et al. 1980)
80
Hanle hot spring
(Fontes et al. 1980)
Dora (#35)
60
Agna
Asal hot spring
(Fontes et al. 1980)
40
Djibouti groundwater
(Fontes et al. 1980; Verhagen et al. 1991)
14
M.E.R.
C(DIC) (% mC)
rift-escarpment CO2
(Bretzler et al. 2011)
Sakalol water
(Fontes et al. 1980; Verhagen et al. 1991)
20
wadi-recharged aquifers
(Fontes et al. 1980; Awaleh et al. 2015)
B
-20
13
local
carbonate
minerals
+8 > δ C > -5 ‰ (Gasse & Fontes 1989)
mantle CO2
-15
-10
-5
0
0
13
δ C(DIC) (‰ vs V-PDB)
13
-2 > δ C > -7 ‰ (Fouillac et al. 1989)
14
Fig. 17. Stable isotope carbon δ13C(DIC) in relationship with carbonate alkalinity concentration as HCO−
C (B), respectively. Colored and open diamonds represent Sakalol waters
3 (A) and
from this study (symbols as in Fig. 16) and the literature, respectively (Fontes et al., 1980; Verhagen et al., 1991). In (A), δ13C values for local carbonate minerals are also shown and plotted
for HCO−
3 ≅ 60 mg/l resulting from water, CaCO3 and logPCO2 = −3 equilibrium at 50 °C (calculated by the PHREEQCI program); dashed lines enclose non-thermal groundwater silicate
13
which has shown high HCO−
3 and low soil δ C(DIC) due to silicate interaction in open-system condition and CO2 from soil (Clark, 2015; Awaleh et al., 2016), whereas arrows depict
possible mixing trends. Red curves depict the trend of the hot springs from the African Rift (Darling et al., 1996), with a CO2 contribution evolving from rift-escarpment to deep mantle
(Bretzler et al., 2011). Symbols and fields are the same as in (B) if available.
similar way, the Abhe and diluted Oued-Kalou hot spring waters, both
interacting with Stratoid basalts, plot near to the evolutive model traced
for the SHGF (Fig. 16). This is in agreement with the chemical and isotope composition, in particular in terms of equilibration with rocks
and the distinction between meteoric and seawater recharge, respectively. On the other hand, the Dora borehole non thermal groundwater
has the most radiogenic value (0.70626) of the study area (Table 2;
Fig. 16), which is within the range of the Sr isotope value
(0.70556–0.70694) of Djibouti volcanic aquifers groundwater (Awaleh
et al., 2016). Therefore, the high 87Sr/86Sr ratios observed for the Dora
borehole groundwater was probably mainly related to the recharge of
the Dora volcanic aquifer by an alluvial-wadi aquifer as observed for
most of the Djibouti volcanic aquifers (Awaleh et al., 2016). C2 and C3
thermal waters plot below the evolutive trend proposed for the SHGF
in the 87Sr/86Sr vs. 1/Sr plot (Fig. 16). This may be due to the nonconservative behavior of dissolved Sr, such as sorption, ion exchange,
or dissolution/precipitation reactions (Bretzler et al., 2011). For example, a diffused presence of hydrothermal calcite with a high Sr concentration in the study area could be an explanation.
The isotope composition of dissolved inorganic carbon (13C/12C and
14
C) of the SHGF waters is interesting when compared with literature
data for Djibouti (Awaleh et al., 2016; Verhagen et al., 1991; Fontes
et al., 1980) and M.E.R. (Bretzler et al., 2011; Craig et al., 1977) waters.
The δ13C(DIC) data from this study is generally more positive than previously published data from the same area (Fig. 17a). This is in accordance with the suggestion of an influx of magmatic CO2, in particular
for C3 waters with a mean of − 4.3 ± 2‰, which is comparable with
the mantle value of − 4 ± 2 (Bretzler et al., 2011; Darling, 2004;
Oppenheimer et al., 2014). Moreover, it should be noted that the
SHGF waters show a low alkalinity value, more compatible with geothermal waters from the Afar such as Allallobeda waters (Craig et al.,
1977) rather than M.E.R. waters (Bretzler et al., 2011; Craig et al.,
1977; Darling et al., 1996). In a similar way, δ13C(CO2) in the Asal geothermal wells waters showed values of −2.6 to −3.1 with null carbonate alkalinity (D'Amore et al., 1998). Unexpectedly, C4 water shows
quite a low value for δ13C(DIC) = − 9.74 ‰, probably due to a reequilibration at surface conditions after degassing, as testified by the
calculated near atmospheric logPCO2 value of − 3.4. Its calculated
δ13C(CO2) of −14.4‰ is compatible with a contamination by soil CO2,
as is more evident for the non-thermal well #36 water (−18.9‰) and
previous-published data from the SHGF and Abhe (Fig. 17a; Fontes
et al., 1980). Similar DIC value was observed for Dora non thermal
water, which falls near to typical wadi-recharged water (Fig. 17a;
Verhagen et al., 1991).
In comparison to published 14C data for Djibouti thermal and nonthermal groundwater, the 14C content of the SHGF waters is one of the
48
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
lowest observed in the country (25 ± 5% pmc), which further confirms
the influx of 14C-free mantle CO2 having δ13C value between
−2 ÷ −6‰ (Fig. 17b). The 13C-14C plot (Fig. 17b) also reveals the possible effect of 13C enrichment due to (re)equilibration with local calcite,
as also shown in Fig. 17a. Radiocarbon dating after a 14C correction by
Netpath XL gave a median value of 6.8 ± 2.4 ka (see supporting information file), which is quite comparable to the period of Asal-Sakalol hydrothermal connection (Gasse and Fontes, 1989). It is interesting to
note that the age calculated for the most ancient C4 water
(8.6–9.0 ka) is compatible with that predicted for the paleowater detected in the Bara desert (Awaleh et al., 2016), whereas the lower values
detected for Cluster 2 and 3 (1.4 ± 1.0 ka and 3.3 ± 2.0 ka respectively)
agree with the that of the “ancient” salt lake in the Dobi graben
(Williams, 2016).
4.8. Conceptual model of the SHGF: an attempted union of hydrological and
geochemical data
There is a few piezometric data in the region. However, the piezometric map of the region, which has been adapted from the country piezometric map (BGR, 1982), has shown clearly that groundwater
recharged from the Moussa Ali Mountain from the north of SHGF
flows towards the Sakalol-Harralol depressions (Fig. S4, see supporting
information). The piezometric map has shown as well that the regional
aquifer groundwater, recharged from the Ethiopian highlights, flows to
the Sakalol-Harralol depressions (Fig. S4, see supporting information).
Moreover, our geochemical results have shown that Asal Rift fluids
flow towards the SHGF.
On the other hand, it has been evidenced from geochemistry and isotope data that thermal waters in the SHGF result from the mixing of
three groundwater: (i) the regional aquifer groundwater; (ii) the
groundwater recharged from the Moussa Ali Mountain; and (iii) the
groundwater that flow from Hanlé area to the Sakalol-Harralol depressions passing through the Asal Rift.
Furthermore, a recent geophysical study (MT, TDEM, Gravimetry)
has shown the possible existence of two geothermal reservoirs in the
SHGF (CERD, 2016): (i) an intermediate geothermal reservoir located
at about 300–600 m depth (probably at T ≅ 110 °C according to
chalcedony equilibrium of C2–C3 samples), and (ii) a deep geothermal
reservoir located about 2000–2500 m depth (T ≅ 143 °C according to
the temperature inferred from C4 sample).
Based on the groundwater movements towards the SHGF, the results of geochemistry and isotope data, and the results of the recent geophysical study on the SHGF, a conceptual model has been proposed for
the SHGF (Fig. 18). It is of interest to note that groundwater recharged
from the Moussa Ali Mountain can be tapped at about 150–200 m
dept as has been noticed in Dora area in the North of the SHGF. Therefore, it is likely that groundwater recharged from the Moussa Ali Mountain feed the intermediate geothermal reservoir of Sakalol-Harralol
geothermal system (Fig. 18). However, only the feasibility study with
drilling geothermal wells and further isotope analysis of intercepted
fluids can make sure if the groundwater recharged from Moussa Ali
Mountain feed as well the deep geothermal reservoir of the SHGF.
The regional aquifer is recharged mainly from the Ethiopian highlight and end up in Djibouti (Gasse and Fontes, 1989). Furthermore, it
has been shown recently that the regional aquifer feed the geothermal
reservoir of Lake Abhe area, located on the south-west of the SHGF
(Awaleh et al., 2015b). Additionally, the geothermal prefeasibility of
Hanlé area (located between Lake Abhé area and SHGF), with one geothermal well with a depth of about 2025 m, showed clearly that such a
regional aquifer (BFGUR, 1999): i) feeds as well the geothermal reservoir of Hanlé geothermal system; ii) can be tapped at about 200 m
depth. On this basis, it is likely that the regional aquifer may feed the intermediate geothermal reservoir of Sakalol-Harralol system as well as
the deep geothermal reservoir of SHGF (Fig. 18).
Our present study confirmed the prediction of the connection between the Asal Rift fluids and Sakalol-Harallol area (Gasse and Fontes,
1989). Therefore, the Asal geothermal fluids, which go through
paleochannel, are likely to feed the deep geothermal reservoir of the
SHGF. However, only feasibility study with drilling geothermal wells
can make sure if this fluid from Asal geothermal field may feed as well
the intermediate geothermal reservoir of the SHGF. Since the Asal Rift
system is recharged mainly by seawater and in some extend by meteoric water (SanJuan et al., 1990), the age of the Asal Rift fluids that transit
to Sakalol-Harralol depressions (at about 9000 y BP) may be due to a
long transit rather than been fossil water.
SW
NE
?
(2)
(3)
I
(T
110°C) ?
II
(T
143°C)
(1)
Heat source
Fig. 18. A simplified sketch of the conceptual model of the Sakalol-Harralol geothermal system. I: Intermediate geothermal reservoir (temperature inferred by chalcedony equilibrium of
the C2–C3 clusters); II: Deep geothermal reservoir (temperature inferred from the C4 sample). (1) Meteoric waters that transit from Asal local rift to the Sakalol-Harralol depressions;
(2) The deep circulating regional aquifer; (3) Groundwater recharged from Moussa Ali Mountain.
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
The schematic section, depicted in Fig. 18, illustrates the possible
flow patterns of the SHGF. Therefore, the deep geothermal reservoir
may mainly been recharged by the regional aquifer groundwater and
the Asal Rift fluids that transit from the Asal local rift to SakalolHarralol depressions (Fig. 18). On the other hand, the intermediate geothermal reservoir may be feed by the groundwater recharged from the
Moussa Ali Mountain and in some extends by the regional aquifer
(Fig. 18). During their ascent by convection through basalt fractures,
the geothermal fluids from the intermediate and the deep geothermal
reservoirs would mix together in varying degrees and in some extend
with the deep circulating regional aquifer, after an additional mixing
with groundwater recharged from Moussa Ali Mountain.. It is of interest
to note that those groundwaters have similar residence time than those
from basalt aquifers encountered in the republic of Djibouti (Awaleh
et al., 2016). The low mineralization of C1 thermal water may be explained by the fact that the groundwater recharged from Moussa Ali
Mountain is to some extend mixed with fresh shallow water during
its ascent.
4.9. Prediction of the scaling
All waters are undersaturated with respect to anhydrite, gypsum,
fluorite, halite and slightly oversaturated in calcite (see supporting informations). In particular, water groups show a progressive increase of
the mean sulfate minerals saturation indices, with no significant distinction between anhydrite and gypsum: C1 = −2.4 ± 0.2; C2 = −1.9 ±
0.2; C3 = − 1.3 ± 0.2. The Galiceela samples (C4) gave the highest
values: SIgypsum = −0.69 and SIanhydrite = −0.39.
The prediction of scaling tendencies of geothermal waters is important in evaluating the production characteristics of geothermal aquifers
and for taking necessary precautions to prevent or control scale formation (Tarcan, 2005). In other words, the saturation indices (log Q/K) of
some carbonate, sulfate, and silica minerals may help one to estimate
which ones of these minerals may precipitate during the extraction
and utilization of the geothermal fluids. Therefore, an assessment of
scaling tendencies may involve the calculation of the saturation state
of these scale-forming minerals. The saturation indices of these latter
in the SHGF thermal waters were computed as a function of temperature. The carbonate minerals (calcite, aragonite, and dolomite) were
found to be oversaturated at 50–300 °C in almost all thermal waters in
the SHGF with the exception of the hot springs Galiceela water (sample
# 90) (Fig. S3, see supporting information). Therefore, carbonate minerals will be most likely to precipitate as scales from the thermal waters
of SHGF during future geothermal production. Indeed, the precipitation
of carbonate minerals from geothermal spring waters in the SHGF is
best illustrated by the occurrence of hydrothermal calcite edifices, several meters high, in the vicinity of still active thermal springs (Fontes
et al., 1979). This formation process is probably the same as that in
Lake Abhe (Fontes and Pouchan, 1975; Fontes et al., 1980; Dekov
et al., 2014).
It is of interest to note that silicate scaling was observed in Asal geothermal wells along with sulphides precipitations (Aquater, 1989;
D'Amore et al., 1998). Indeed, fluids produced from Asal geothermal
wells have shown strong scaling tendency related mainly to sphalerite
and galena precipitation in the production casing and silica precipitation in wellhead (Aquater, 1989; D'Amore et al., 1998). However, carbonate precipitation was not reported for Asal geothermal wells
(Aquater, 1989).
5. Conclusions
Hot springs in the Sakalol-Harralol geothermal field were found to
be aligned NW-SE along the main faults. Four groups of waters (Clusters
1–4) were defined using HCA and PCA and an excellent correspondence
was found with their geographical location. Alkaline thermal waters in
the Sakalol-Harralol geothermal field were found to be high TDS Na-Cl
49
type waters, while non-thermal groundwater or the freshest thermal
waters are mostly Na-Cl-HCO3-SO4 waters. The increase of salinity
from Cluster 1 to Cluster 4 is due to a combination of contributions
from evaporites, mainly evident in Cluster 3 and typical of the “passive”
graben setting in the Afar, and a recycled seawater component of the
geothermal water of the “active” Asal rift, mainly evident in Cluster 4
(approximatively 7%). Those contributions are obliterated by progressive equilibration with Stratoid basalt; however, the seawater imprint
is particularly evident from the elevated values of boron isotopes and
the relatively high chloride concentration in comparison with more diluted samples. On the contrary, and as expected in rift setting (Shanks
et al., 1981; Awaleh et al., 2015a), no trace of seawater was revealed
by the isotope composition of dissolved sulfate. Instead, a dominant
contribution of sulfate derived from magmatic SO2 disproportionation,
which may take place locally in a different form (i.e. included in basalt
rock, recycled in sedimentary evaporate, fumarolic condensate) was
proven. Different contributions in term of water recharge spread the
samples in terms of water isotope composition. Cluster 1, Cluster 4
and Clusters 2–3 are probably fed with meteoric water from Moussa
Ali Mountain, meteoric water that transit from Asal local rift to
Sakalol-Harralol depression and meteoric water from the deep circulating regional aquifer respectively. This latter may be the result of different, mixed contributions.
Locally, the deep geothermal reservoir may be recharged mainly
by the regional aquifer groundwater and the Asal Rift fluids that
transit to Sakalol-Harralol depressions. Whereas, the intermediate
geothermal reservoir may be feed by the groundwater recharged
from the Moussa Ali Mountain and in some extends the regional
aquifer. During their ascent, the geothermal fluids from the intermediate and the deep geothermal reservoirs would mix together and in
some extend with the deep circulating regional aquifer, with an additional mixing with groundwater recharged from Moussa Ali
Mountain.
Different geothermometrical approaches give a temperature range
estimates for the deep geothermal reservoir of approximatively
120–160 °C, with a mean deep temperature of 143 °C. According to previous hydrological studies, it could be also likely the presence of a shallow aquifer at 110 °C. Common features in terms of deep temperature,
chemical and water-isotope composition were revealed between the
Sakalol-Harralol deep geothermal water and the Hanlé, Abhe, Dobi graben and diluted shallow aquifer waters of the Asal geothermal field.
Therefore, the so called regional aquifer may have a wider extension
than previously assumed. Finally, the results of geochemical studies
and field observation suggest that the most likely scales to be precipitated during the extraction of thermal water in the study area are carbonate minerals. Apart from being a step forward in the characterization of
the geothermal resource of the Sakalol area and therefore bringing the
key knowledge for an utilization of this renewable energy, this study
will serve as a valuable example of a large multi-tracer isotope and geochemical characterization of a complex geothermal system.
Acknowledgments
This research work was financially supported by the Centre d'Etudes
et de Recherche de Djibouti (CERD). We are grateful to Abdillahi Elmi
Adaneh, Ismael Said Ismael and Sikieh Abdillahi Elmi for their assistance
in the field works. We would like to thank the Djibouti National Army
for their assistance in logistic. We would also like to thank two anonymous reviewers for their constructive comments that improved the
manuscript.
Appendix A. Supplementary data
Supplementary data to this article can be found online at http://dx.
doi.org/10.1016/j.jvolgeores.2016.11.008.
50
M.O. Awaleh et al. / Journal of Volcanology and Geothermal Research 331 (2017) 26–52
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