ORIGINAL ARTICLE

Hydrogeochemical evaluation of groundwater in the lowerOffin basin, Ghana

Benony K. Kortatsi Æ Collins K. Tay ÆGeophrey Anornu Æ Ebenezer Hayford ÆGrace A. Dartey

Received: 3 February 2007 / Accepted: 17 April 2007 / Published online: 17 May 2007

� Springer-Verlag 2007

Abstract Alumino-silicate mineral dissolution, cation

exchange, reductive dissolution of hematite and goethite,

oxidation of pyrite and arsenopyrite are processes that

influence groundwater quality in the Offin Basin. The main

aim of this study was to characterise groundwater and

delineate relevant water–rock interactions that control the

evolution of water quality in Offin Basin, a major gold

mining area in Ghana. Boreholes, dug wells, springs and

mine drainage samples were analysed for major ions, minor

and trace elements. Major ion study results show that the

groundwater is, principally, Ca–Mg–HCO3 or Na–Mg–Ca–

HCO3 in character, mildly acidic and low in conductivity.

Groundwater acidification is principally due to natural

biogeochemical processes. Though acidic, the groundwater

has positive acid neutralising potential provided by the

dissolution of alumino-silicates and mafic rocks. Trace

elements’ loading (except arsenic and iron) of groundwater

is generally low. Reductive dissolution of iron minerals in

the presence of organic matter is responsible for high-iron

concentration in areas underlain by granitoids. Elsewhere

pyrite and arsenopyrite oxidation is the plausible process

for iron and arsenic mobilisation. Approximately 19 and

46% of the boreholes have arsenic and iron concentrations

exceeding the WHO’s (Guidelines for drinking water

quality. Final task group meeting. WHO Press, World

Health Organization, Geneva, 2004) maximum acceptable

limits of 10 lg l–1 and 0.3 mg l–1, for drinking water.

Keywords Cation exchange � Ghana � Offin basin �Silicate weathering � Pyrite and arsenopyrite � Oxidation

Introduction

The Lower Offin Basin is a major gold-mining area in

Ghana. This area receives high-annual rainfall and as a

result is drained by numerous springs and streams (Dickson

and Benneh 1980). These streams were the sources of

drinking water to some of the communities in the Basin and

also act as pathways through which weathered materials

and mine effluents were transported into the sea via the

major river, River Pra. Due to the discharge of mine

effluents into the rivers and streams by the major mining

companies and the activities of small-scale miners, most of

these surface water bodies have become polluted and

unsuitable for drinking water supply. Groundwater has

therefore become the principal and sometimes the only

source of drinking water in the area. Consequently, a large

number of boreholes and hand-dug wells have been drilled

in the area and a lot more are being considered.

Groundwater in hard-rock aquifers, particularly in min-

ing areas, is known to be vulnerable to quality problems that

may have serious impact on human health. The rocks are

often carbonate-deficient and give rise to poorly buffered

water (Smedley et al. 1995). Also in gold and base metal

mining areas such as the Lower Offin Basin, inorganic and

microbial-induced sulphide oxidations are a common phe-

nomenon. This often leads to low pH groundwater that

B. K. Kortatsi (&) � C. K. Tay � G. A. Dartey

CSIR-Water Research Institute,

P. O. Box M32, Accra, Ghana

e-mail: kortatsi@hotmail.com

G. Anornu

Civil Engineering Department, KNUST,

Kumasi, Ghana

E. Hayford

Geology Department, University of Ghana,

Legon, Accra, Ghana

123

Environ Geol (2008) 53:1651–1662

DOI 10.1007/s00254-007-0772-0

might encourage dissolution of trace elements such as Al,

As, Be, Cr, Mn and Fe from most host rocks in dangerous

proportions into the groundwater system if they are present

in the rock matrix. Wilson and Hawkins (1978) observed

concentrations of As between 240 and 1,200 lg l–1 in the

Fairbanks area, Alaska. Williams and Smith (1994) reported

that acid waters draining in a gold-mining area in Zimba-

bwe had As concentration of up to 72 mg l–1. Therefore,

basing the water supply of the study area on groundwater

without adequate information on quality or the chemistry of

groundwater can be dangerous.

A few studies pertaining to groundwater quality were

carried out in the Offin Basin. These were specifically in

connection with arsenic mobilisation and its health impli-

cations on rural water supply and include Amasa (1975),

Smedley et al. (1995), Smedley et al. (1996) and Smedley

(1996).

In all these studies, only little emphasis has been placed

on either the natural processes that govern the chemical

composition of the groundwater or the anthropogenic fac-

ets, which have some effects on it. The main objective of

this paper, therefore, is to use existing chemical data to

characterise the groundwater and to delineate the relevant

water–rock interactions and anthropogenic factors that

control the water quality in the study area.

Study area

The Lower Offin Basin lies between latitudes 5.8�N and

6.5�N and longitudes 1.3�W and 2.2�W (Fig. 1). This area

receives copious rainfall between March and October and

its mean annual value is approximately 1,600 mm. The

diurnal temperature is between 20 and 30�C while the

potential evapotranspiration is about 1,260 mm year–1 al-

though it does not exceed rainfall for most part of the rainy

season. The Lower Offin Basin is covered with the moist

semi-deciduous forest although much of the vegetation has

been removed for cocoa and food crop production. Tec-

tonic lines of weakness and Eburnian orogenic episodes

have facilitated extensive and intensive regolith and ravine

development. Consequently many perennial streams and

rivers (e.g. Ofin, Oda and Gymi) drain the area.

The geology of the area (Fig. 2.) consists mainly of the

Proterozoic Birimian that comprises of the Lower Birimian

(metasediment rocks) and the Upper Birimian (metavol-

canic rocks). The lower Birimian consists mainly of schist,

phyllites, metagreywacke and granites while the Upper

Birimian is made up mainly of a bedded group of green

lavas (greenstones), tuffs and sediments with minor bands

of phyllite (Kesse 1985). Minor rock types include Basic

Intrusives and Tarkwaian Formation.

Land use

Land use in the Offin Basin is mainly of gold mining and

farming (cocoa and food crop).

Gold mining in the area are of two types, large-scale

mining and small-scale mining. Large-scale mining is

conducted by heap leach technique or roasting of ore.

Oxidised ores derived from sulphides principally arseno-

Fig. 1 Location map of the

Offin Basin showing the

sampling points

1652 Environ Geol (2008) 53:1651–1662

123

pyrites, realgar, orpiment and pyrites in the weathered

zones are heap leached by cyanidation. Paleoplacer (free

milling ore) are mined from deep zones crushed, milled

and cyanided. Small-scale mining involves extracting gold

from alluvial soils mainly from stream floors by mercury

amalgamation.

Cocoa farming is the main cash crop farming in the

study area. Large acreages of flora or virgin forest were

removed and replaced with cocoa farms. Additionally, food

crops such as plantain, cassava, cocoyams etc. are pro-

duced alongside cocoa for subsistence. Due to the fact that

agro chemicals are expensive and often above the reach of

most the peasant farmers, their use in the study area is

limited.

Materials and methods

Sample collection

Samples were collected from 148 water points (boreholes,

hand-dug wells and springs) and 2 samples from mine

drainage in December 2004 during the dry period. Sam-

pling protocols described by Claasen (1982) Barcelona

et al. (1985) and Gale and Robins (1989) were strictly

observed during sample collection. Samples were collected

using 4-1 acid-washed polypropylene containers. Each

sample was immediately filtered on site through 0.45 lm

filters on acetate cellulose. Filtrate for metals analyses were

transferred into 100-cm3 polyethylene bottles and imme-

diately acidified to pH < 2 by the addition of MerckTM

ultrapure nitric acid (5 ml 6N HNO3). Samples for anions

analyses were collected into 250-cm3 polyethylene bottles

without preservation. All the samples were stored in an ice

chest at a temperature of <4�C and later, transferred to the

laboratory and stored in a refrigerator at a temperature of

<4�C until analysed (within 1 week).

Field measurements

In the field, temperature, pH and electrical conductivity

measurements were carried out using WTW-Multiline P4

Universal Meter in flow-through cell attached in line to the

borehole pump outlet. Before taking readings, pumping

was carried out until the meter readings were stable for

each parameter. Since the boreholes were already being

used, the pumping time before stable meter readings were

achieved was, on average, 5 min. The pH electrode was

calibrated against pH 4, 7 and or 10 buffers (Wood 1981).

Alkalinity (as HCO3–) was determined by field titration with

1.6 N H2SO4 to pH ~4.5 using HACH Digital multi

Sampler Model 1690

Laboratory measurements

Chemical analysis for anions (Bicarbonate, Chloride, Sul-

phate and nitrate was performed at the Ecological labora-

tory at the University of Ghana using Dionex DX-120 ion

N

G H A N A

I N S E T

0 10 20 30 Km.

Fig. 2 Geological map of

the Offin Basin with the

sampling points

Environ Geol (2008) 53:1651–1662 1653

123

chromatography. Detection limits were 2 lmol 1–1 for

SO42–, NO3

–, HCO3 and Cl–. Trace element analyses were

performed on the acidified samples using the Unicam 969

Atomic Absorption Spectrometer (AAS). An ionic balance

was computed and used as a basis for checking analytical

results.

The ionic balances for the analyses varied from 5.9 to

10.1%, though more than 85% of the samples had an ionic

balance within ±5%. An ionic balance outside ±5% was

largely associated with samples with very low total dis-

solved solids.

Results and discussion

A statistical summary of chemical parameters measured in

the 150-groundwater samples is presented in Table 1. The

pH of the groundwater is typically low and varies from 3.9

to 7.0 with mean and median values of 5.8 and 5.7.

However, the majority of the samples (>95%) fall within

natural water pH range 4.5–7.0 (Langmuir 1997). Electrical

conductivity (EC) values are generally low. Minimum and

maximum values are 37.3 and 956.0 lS cm–1 with mean

and median values of 185 and 138, respectively.

Table 1 Statistical summary of

the parameters determined in

the 148 groundwater samples

Temp. in �C, pH in pH-units,

EC in lS cm–1,Ca to Al in

mg l–1, AS to Zn in lg l–1

Detection

limit

Min Max Mean Median STD WHO (2004)

Guideline limit

Temp 22.1 29.3 25.8 26.0 1.3

pH 3.9 7.0 5.9 5.9 0.6

Cond 37.3 956.0 168.2 124.0 139.3

Ca 0.01 0.3 87.3 11.6 6.6 15.2

Mg 0.01 0.1 77.1 5.2 3.6 8.9

Na 0.1 2.2 71.3 12.8 10.8 8.4 200

K 0.01 0.0 19.2 1.6 0.8 2.5

HCO3 0.1 0.1 258.0 67.0 51.5 55.7

SO4 0.2 0.1 507.0 10.5 0.8 57.4 250

Cl 0.1 2.0 43.0 8.1 6.2 6.6 250

NO3 0.2 0.2 21.9 1.4 0.2 3.3 50

P 0.1 <0.1 0.5 0.1 <0.1

Si 0.1 3.6 39.4 17.0 14.0 9.7

F 0.05 0.01 11.10 0.20 0.07 1.08 1.5

Fe(total) 0.003 <0.003 16.800 1.276 0.233 2.529 0.3

Al 0.1 0.02 2.39 0.10 0.02 0.34 0.2

As 0.05 0.00 64.00 8.79 1.55 15.65 10

B 0.5 6.00 118.00 8.52 6.00 11.31 300

Ba 0.04 1.00 502.00 40.82 17.50 75.63 700

Cd 0.02 <0.02 7.2 0.1 0.02 0.2 3

Cr 0.2 <0.2 205.4 2.7 2.7 1.2 50

Cu 0.04 4.00 62.00 8.12 4.00 9.59 2,000

Hg 0.05 <0.05 3.85 0.18 0.05 1.8 1

Mo 0.03 <0.03 4.6 0.1 0.03 0.1 70

Mn 0.02 1.00 1020.00 116.97 52.05 155.75 500

Ni 0.06 <0.06 746.1 9.8 7.28 10.3 20

Pb 0.03 0.56 167.5 2.2 1.33 3.2 10

Rb 0.01 <0.01 47.4 0.6 0.33 0.9

Sb 0.025 <0.05 7.2 0.1 0.05 0.3 5

Se 0.5 <0.5 254.6 3.4 1 4.2 10

Sn 0.05 <0.05 4.1 <0.06 0.05 0.3

Sr 0.1 4.00 546.00 116.41 63.00 123.89

Th 0.02 <0.02 1.4 <0.02 0.02 0.

U 0.01 <0.01 0.9 <0.01 0.01 0.07

V 0.02 <0.02 0.88 0.33 0.18 0.26

Zn 0.1 4.00 1500.00 47.28 12.00 151.91 3,000

1654 Environ Geol (2008) 53:1651–1662

123

Major cations (Ca2+, Mg2+, Na+ and K+) are also gen-

erally low. Comparison of their concentrations is illustrated

in the box plot (Fig. 3). Na+ is relatively the dominant

cation with median value of 0.5 mmol/l and a narrow

spread. HCO3– is by far the dominant anion with median

value of 0.9 mmol/l and forms approximately 82% of the

anion content. SO42– has a low median concentration but is

largely spread. Figure 3 illustrates that large percentages of

contribution to salinity as depicted in Fig. 4 are from

HCO3–, Na+ and Ca2+. The spatial distribution of salinity is

illustrated with the contour map of electrical conductivity

of the groundwater samples (Fig. 4) since conductivity is

sensitive to variations in dissolved solids (mostly mineral

salts). The degree to which the salts dissociate into ions, the

Na Ca Mg Cl HCO3 SO4

Parameters

0.001

0.010

0.100

1.000

10.000

100.000

Con

cent

ratio

ns (

mm

ol/l)

Box and Whisker Plot

Legend

Max.

75 percentile

Median

25 percentile

Min.

Fig. 3 Box and whisker plot of

major chemical constituent in

the groundwater of the study

area

-1.9 -1.8 -1.7 -1.6

6.1

6.2

6.3

6.4Atobiase

Adansi

Odumase

Boni

Manfo

Antoakurom Asaman

Akrofuom

Mensakurom0

100

200

300

400

500

600

700

800

0 10 20 30 Km

Conductivity (uS/cm)

Fig. 4 Contour map of

electrical conductivity

distribution in the groundwater

of the Offin Basin

Environ Geol (2008) 53:1651–1662 1655

123

amount of electrical charge on each ion, mobility and the

temperature of the solution, all have influence on conduc-

tivity (UNESCO/WHO/UNEP 1996). In other words,

conductivity is a linear function of total dissolved solids

and for that matter the mineral salt content of the

groundwater (Hem 1970; Freeze and Cherry 1979). The

conductivity contours show a trend that depicts localised

flow cells. Nonetheless, conductivity values are generally

low in the mid portion, relatively moderately high in the

north and high in the southwest.

Water types

Figure 5 shows Piper Trilinear Plot (Piper 1944) of major

ions concentrations of water samples in the study area. Two

principal hydrochemical water types have been delineated.

These are Ca–Mg–HCO3-type water where the chemical

properties of the water are dominated by alkaline earths

and weak acids (Karanth 1994). This water type occupies

the section of the diamond shape designated as I. The Na–

Mg–Ca–HCO3 is the second water type. In this water type,

no particular cation predominates while HCO3 is the main

anion. It occupies the section of the diamond shape des-

ignated by II. Other minor water types include Na-HCO3

water type (III section of the diamond shape), Na–Cl or

Na–HCO3–Cl occupying the section designated by IV and

Ca–Mg–SO4 water (V field) of the diamond shape. The

principal water types and Na–HCO3 water type depict

rock–water interaction involving the dissolution of car-

bonates and feldspars by the recharging groundwater

within the relatively permeable weathered zone above the

underlying rocks. The rock types of the study area, as

stated earlier in the sections on geology, consist mainly of

calc-chlorite schist, carbonate–albite dolerite, biotite and

muscovite–biotite granite and adamellite. Garrels and

Mackenzie (1967) pointed out that meteoric water dis-

solving Na from Na-bearing silicates would produce

NaHCO3 water type. Therefore, albite dissolution would

probably have produced NaHCO3 type. Na–Cl or Na–

HCO3–Cl water type has low conductivity suggesting a

meteoric origin. A few of the boreholes and mine drainage

samples plot close to the top corner of the diamond shape

and towards the SO4 apex in the anion field while in the

cation field, plot almost in the middle along the Ca, Mg

axis indicating that this water type evolved from pyrite or

arsenopyrite oxidation.

Chemical controls

The processes responsible for the chemical character of a

particular water type can be deduced using chemical

parameter ratios, which are based on simplistic mass bal-

ance approach to water quality data (Hounslow 1995).

The Cl/P

anions ratios of the groundwater samples are

generally low and in the range (0.0–0.7) with median value

of 0.2 (Table 2). On the contrary, the HCO3/P

anions

Fig. 5 Piper plot of water

samples from the Offin Basin

1656 Environ Geol (2008) 53:1651–1662

123

ratios are high with median 0.8 suggesting rock weath-

ering, principally silicate or carbonate weathering as has

been observed on the Piper plot. The occurrence of Na–

HCO3 water also suggests the possibility of ion exchange.

Thus, the water chemistry is probably controlled by car-

bonate or silicate dissolution and cation exchange. The

dominance of silicate or carbonate-weathering process can

be deduced from HCO3/SiO2 molar ratio. If the ratio is

greater than ten, which implies HCO3 >> SiO2, carbonate

weathering predominates. However, if the ratio is low and

in fact less than five, then silicate weathering predominates

(Hounslow 1995). Apart from a few of the water samples

that showed high ratios (HCO3/SiO2 > 10), most of the

samples had low to moderate ratios (median value = 3.6)

suggesting silicate weathering as a more dominant chemi-

cal process influencing the chemical character of ground-

water in the study area than carbonate weathering. The low

Cl/P

anions ratio of the groundwater in an area remote

from the sea suggests that the Cl– was derived mainly from

atmospheric sources and, therefore, the Cl– would be

present in the groundwater as sodium chloride (Karanth

1994; Hounslow 1995). Thus the Na+: Cl– molar ratio

would be expected to be 1:1. Figure 6 is a plot of Na+

against Cl–; it can be observed that most of the samples

plotted above the 1:1 line indicating excess of Na+ over

Cl–. Excess Na+ can be derived from sources that include

plagioclase (albite) dissolution as per Eq. (1) or cation

exchange (natural water softening) as shown by Eq. (2)

2NaAlSi3O8 þ 2CO2 þ 11H2O! Al2Si2O5ðOHÞ4þ 2Naþ þ 2HCO�3 þ 4H4SiO4 ð1Þ

0:5 Ca2þ þ Na� clay! 0:5 Ca� clayþ Naþ: ð2Þ

Cation exchange

In order to investigate the occurrence of cation exchange

reaction in the groundwater, Ca + Mg – (HCO3 + SO4)

(meq) was plotted against Na–Cl (meq.) (Fig. 7). Since

calcite, dolomite, gypsum and anhydrite are the most likely

additional sources that Ca2+ and Mg2+ could enter the

groundwater apart from cation exchange, in plotting this

diagram, possible contributions of Ca2+ and Mg2+ from

calcite, dolomite, gypsum and anhydrite dissolution to

lithogenic Ca2+ and Mg2+ in the groundwater were ac-

counted for by subtracting the equivalent concentrations of

HCO3 and SO4 (Nkotagu 1996; McLean and Jankowski

2000). Similarly to account for lithologenic Na+ available

for exchange, it was assumed that Na+ contribution from

meteoric origin would be balanced by equivalent concen-

tration of Cl– and, therefore, equivalent Cl– concentration

was subtracted from that of Na (Nkotagu 1996; McLean

and Jankowski 2000). A large percentage (approxi-

mately70%) of the water sample plot away from the zero of

the x-axis and fall approximately along a line with gradient

of –1. Water undergoing ion exchange plots along a line

with a slope of –1 in Ca + Mg – (HCO3 + SO4) plot field

(Jankowski et al. 1998). Thus, ion exchange is a significant

process through which Na+ partly enters the groundwater

system.

Silicate weathering

Si concentration is generally low with median value of

14.0 mg l–1. Possible sources of silica in the groundwater

include weathering of quartz, feldspars and ferromagnesian

Table 2 Summary statistics of saturation indices (si) of common

mineral phases in the rocks; ratios of some chemical constituents

Parameter Min Max Mean Median Stdev

pH 3.9 7.0 5.9 5.9 0.6

si_Calcite –6.8 1.4 –2.9 –2.9 1.4

si_Dolomite –13.6 1.5 –5.8 –5.5 2.6

si_Gypsum –6.4 –0.1 –4.3 –4.4 1.2

si_Anhydrite –6.7 –0.3 –4.5 –4.6 1.2

si_Goethite –3.5 9.1 5.6 5.9 2.7

si_Hematite –5.1 20.3 13.2 13.9 5.3

si_Fe(OH)3(a) –9.5 3.2 –0.3 0.0 2.7

si_SiO2(a) –1.5 –0.5 –0.9 –1.0 0.3

si_Quartz –0.3 0.8 0.3 0.3 0.3

SO42– (mmol l–1) 0.001 5.28 0.11 0.01 0.59

HCO3/H4SiO4 0.6 28.0 4.4 3.6 4.0

HCO3/P

anions 0.04 0.96 0.68 0.75 0.23

Cl/P

anions 0.02 0.57 0.18 0.16 0.11

0

0.5

1

1.5

2

2.5

3

3.5

0 0.2 0.4 0.6 0.8 1 1.2 1.4

Cl (meq/l)

Na

(meq

/1)

deep well (>25 m)Shallow well (<15 m)SpringLinear (01:01)

egnahcxe noitaC

gnirehtaew etaciliS

Reverse ion-exchange

Legend

Fig. 6 Plot of Na versus Cl for groundwater samples from the study

area

Environ Geol (2008) 53:1651–1662 1657

123

silicates. Figure 8 is the stability diagram of Ca-silicates

and Na-silicates and their weathering products. This dia-

gram was drawn with the assumption that aluminium was

preserved in the weathering products (Appelo and Postma

1999). End member compositions were also assumed using

equilibrium relationships of Tardy (1971) for standard

temperature (25�C) and pressure (1 atmosphere), which

approximately reflected conditions in the groundwater.

Constituents’ activities were computed using Phreeqc for

Windows Version 2, 12.04 (Parkhurst and Appelo 1999).

Figure 8 illustrates that the groundwater is under saturated

with respect to amorphous silica while large percentage of

the water samples are either saturated or supersaturated

with respect to quartz quite similar to what has been ob-

served in most natural waters (Hem 1992). Silica if avail-

able may continue to dissolve until amorphous silica

becomes saturated. Thus, amorphous silica is apparently

the control on the dissolution of silica in the groundwater.

Consistent with climatic conditions, the entire groundwater

samples plot in the kaolinite-stability field indicating that

kaolinite is the most stable secondary silicate mineral phase

for the groundwater system.

Saturation indices

The state of saturation of the groundwater with respect to the

major carbonate mineral species was assessed in order to

investigate the thermodynamic controls on the composition

of the groundwater and the approximate degree to which the

groundwater has equilibrated with the various carbonate

mineral phases since HCO3– is the dominant anion. The sat-

uration indices were determined using the hydro geochemi-

cal equilibrium model, Phreeqc for Windows (Parkhurst and

Appelo 1999). The saturation index (SI) of a given mineral is

defined in Eq. (3) (Garrels and Mackenzie 1967)

SI ¼ log10ðIAP=KspÞ ð3Þ

where IAP is the ion activity product, Ksp is the solubility

product at a given temperature (the thermodynamic equi-

y = -1.09x + 0.0038R2 = 0.63

-2.5

-2

-1.5

-1

-0.5

0

0.5

1

-0.5 0 0.5 1 1.5 2 2.5

Na-Cl (meq/l)

)l/qem( })4

OS+3O

CH(- g

M+ aC{

Fig. 7 Bivariate plot of

Ca + Mg – HCO3–SO4 versus

Na–Cl for the groundwater

samples from the Offin Basin

20.0

18.0

16.0

14.0

12.0

10.0

8.0

6.0

4.0

2.0

0.0-7.0 -6.0 -5.0 -4.0 -3.0 -2.0 -1.0 -0.0

GIBBSITE

KAOLINITE

Na-MONTMORILLONITE

ALBITE

Log[H4SiO4]

Log

([N

a+]/

[H+

])ANORTHITE

Ca-MONTMORILLONITE

Log

([C

a²+

]/[H

+]²

)

L E G E N D

Mine DrainageSpringDeep WellShallow Well

Sodium - Silica system

Calcium - Silica system

Fig. 8 The stability of anorthite

and albite and their possible

weathering products gibbsite,

kaolinite and Ca-

montmorillonite with respect to

the groundwater in the Offin

Basin

1658 Environ Geol (2008) 53:1651–1662

123

librium constant adjusted to the temperature of the given

sample).

The thermodynamic data used in this computation were

those contained in the default database of the ‘Phreeqc for

Windows’. Supersaturation (SI > 0) indicates that precip-

itation is thermodynamically favourable in spite of the fact

that slow rates of reaction can inhibit precipitation. On the

other hand, undersaturation (SI < 0) signifies that dissolu-

tion is favoured. The summary statistics of saturated indi-

ces of some of the common mineral phases is presented in

Table 2, which clearly shows that the groundwater is

generally undersaturated with respect to most of the car-

bonate phases. However, petrographic evidence indicates

that carbonates exist in the rock matrices, thus the under-

saturation may be attributable to very short residence time

of the groundwater in the aquifer or the carbonate species

may exist in other forms, which are not readily soluble in

water. The groundwater is also undersaturated with respect

to gypsum and anhydrite due plausibly to the absence of

these mineral phases from the rock matrix. It is also uni-

formly supersaturated with respect to quarts but undersat-

urated with respect to amorphous silica (Table 2)

suggesting that amorphous silica possibly controls the sil-

ica concentration in the groundwater.

Minor and trace elements

The analytical results did not indicate high levels of minor

and trace metal loading expected from mining area. None-

theless, arsenic and iron showed concentrations significantly

above their detection levels in a few boreholes sampled.

The concentration of arsenic is generally low varying

from <0.05 to 64.00 lg l–1 with a median value of

1.55 lg l–1. Relatively high As concentrations (>5 lg l–1)

occur in the western and the northwestern part of the study

area, where the Birimian rocks contains significant amount

of pyrite and arsenopyrites (Fig. 2). Approximately 19% of

the boreholes have arsenic concentrations slightly to

moderately in excess of the WHO guideline limit of

10 lg l–1 (WHO 2004) suggesting arsenopyrite oxidation

enhanced perhaps by mining activities in the area; arse-

nopyrite oxidation can take place without mining once the

pyrite or arsenopyrite are exposed to air and water as in

weathering reactions. Oxidation of pyrite and arsenopyrite

may be expressed by Eqs. (4) and (5), respectively

FeS2 þ 3:5O2 þ H2O ¼ Fe2þ þ 2SO2�4 þ 2Hþ ð4Þ

FeAsS þ 3:25O2 þ 1:5 H2O

¼ Fe2þ þ AsO3�4 þ SO2�

4 þ 3Hþ ð5Þ

These reactions are supposed to produce high sulphate

concentration, low pH and Fe2+/SO42– molar ratios of 0.5

and 1.0, respectively. However, SO42– concentration in the

boreholes is generally very low, 0.001–5.28 mmol l–1 with

a median value of 0.01 mmol l–1. Additionally Fe2+/SO42–

molar ratios are generally very low (median value = 0.1)

and only a few samples satisfy iron to sulphate (Fe2+/SO42–)

molar ratios of 0.5 and 1.0 for the stoichiometry of pyrite

and arsenopyrite oxidation, respectively. It is possible that

iron may have precipitated as amorphous Fe(OH)3 and

hydrated to form limonite or hematite as in Eq. (6).

FeS2 þ 3:75O2 þ 2:5 H2O ¼ FeOOH þ 2SO2�4 þ 4Hþ:

ð6Þ

This is possibly followed by sulphate reduction while

pH remains relatively moderate because of sediment buf-

fering. Though, no field evidence pointed to sulphate

reduction taking place, mineralogy of the weathered profile

in the Obuasi area within the South Offin Basin has pointed

to the presence of arsenate and ferric hydroxide minerals as

the product arsenopyrite oxidation (Bowell 1994), which

probably act as sink for As and other trace metals. The pe-

pH diagram for the groundwater system at 25�C and 1

atmosphere within the study area is presented in Fig. 9.

Stability constants are taken from Appelo and Postma

(1999). Consistent with field observation, the samples plot

in the Fe2+/Fe(OH)3 (FeOOH) indicating amorphous iron

as the control on Fe concentration in the groundwater.

Nonetheless, boreholes with high iron concentration also

Fig. 9 Stability diagram for iron species in groundwater. Boundaries

valid at 25�C, 1 atmosphere total pressure, total iron activity of

dissolved species 10–6

Environ Geol (2008) 53:1651–1662 1659

123

have Fe2+/SO42– > 1 suggesting that unless SO4

2– reduction

has taken place, but as said earlier has not been evident,

pyrite and/or arsenopyrite oxidation may have occurred but

cannot explain the high iron concentration in the ground-

water particularly at areas underlain by granites and basic

intrusives in the northeast. Furthermore, calculated satu-

rated indices of the iron species goethite, hematite and

magnetite indicate that the waters are highly supersaturated

with respect to these species. Thus the groundwater should

be precipitating these iron minerals to reduce Fe2+/SO42–

ratio. Since this does not seem to happen, Fe2+ could have

been derived from other sources apart from pyrite or

arsenopyrite oxidation. Hematite, scorodite and goethite

and to a limited extent magnetite are some of the iron

species present in the rock matrix of the study area (Kesse

1985; Bowell 1992). Reductive dissolution of these iron

minerals results in the production of Fe2+ that can sustain

Fe2+/SO42– > 1. Hematite dissolution in the presence of

proton (acid) is extremely slow but in the presence of or-

ganic ligands the reaction can be fast (Appelo and Postma

1999). Though in this study the content of organic matter in

the soil zone has not been measured, the upper zone of soils

to the east of the study area is known to be high in organic

content (Bowell 1993). The brown colouration of rivers

and streams throughout the year in Offin basin is a further

indication of the presence of organic matter in the soil

zone. Thus, reductive dissolution of iron minerals in the

presence of organic matter and perhaps a reductant seems

to be a more plausible process through which Fe2+ entered

the groundwater particularly in the low sulphide ore areas

in the north–east where the major rock types are granites

and granodiorites than arsenopyrite or pyrite oxidation.

Additional source of iron could be the leaching from fer-

romagnesian silicates—Biotite, Chlorite, hornblende and

pyroxenes, which are common in the rock types that

underlain the study area.

Groundwater acidification

Total hardness-alkalinity plot is usually used to evaluate

the acidity status of ground and surface waters (Henriksen

1979; von Bromsse 1989; Knutsson 1994; de Caritat 1995).

Figure 10 illustrates the relationship between total hard-

ness and alkalinity of groundwater in the Offin Basin. Most

of the groundwater samples in Fig. 10 plot along the 1:1

line suggesting that acidification is principally due to nat-

ural biogeochemical processes (CO2 generation in the soil

zone through root respiration and decay of organic matter).

The few samples that plot along or above the 2:1 line

probably resulted from sulphide (pyrite or arsenopyrite)

oxidation enhanced by mining. This confirms that acidifi-

cation due to acid mind drainage is limited in extent in the

groundwater.

Groundwater acidification (Aci.) can also be defined as

loss of alkalinity and computed using Eq. (7).

Aci: ¼ 0:93 ðCa� þMg�Þ � 14� Alkþ Al ð7Þ

where all concentrations are expressed in meq l–1

(Henriksen 1980; Henriksen and Kirkhusmo 1986; Caritat

et al. 1998). Similarly, the acid neutralising capacity of the

groundwater is computed using Eq. (8).

ANC ¼ ðCa2þ þMg2þ þ Naþ þ KþÞ� ðNO�3 þ SO2�

4 þ Cl�Þ ð8Þ

where all concentrations are expressed in meq l–1 (Stumm

1992; Caritat 1998).

Figure 11 presents the overview of Aci, ANC and

NANC (Net acid neutralising capacity) of groundwater in

1:1

2:1

0

2

4

6

8

10

12

14

0 1 2 3 4 5 6 7

Alkalinity

ssendraH lato

T

Deep well (>25m)

Shallow well (<15m)

Spring

Legend

Fig. 10 Relationship between total hardness and alkalinity of

groundwater in the Offin Basin

0.00

4.00

8.00

12.00

16.00

20.00

-16.00 -14.00 -12.00 -10.00 -8.00 -6.00 -4.00 -2.00 0.00

Aci (meq/l)

)l/qem(

CN

AN ;

CN

AANCNANC

Fig. 11 Relationship between acid-neutralising capacity (ANC); net-

acid neutralising capacity (NANC) and acidification capacity of

groundwater in the Offin Basin

1660 Environ Geol (2008) 53:1651–1662

123

the Offin Basin. It illustrates low, even negative values for

acidification at almost all the sampling points. ANC,

though low remains positive (in the range 0.1–4.2 meq/l)

signifying that the groundwater still has the potential to

neutralise acid. The high NANC illustrated by the plot is

evocative of the fact that the groundwater has not been

significantly acidified. The acid neutralisation potential is

probably provided by the dissolution of alumino-silicates

and some mafic rocks.

Drinking water quality

Generally, groundwater quality in the Offin Basin is good

for drinking purposes as most of the chemical constituents

are well within the WHO (2004) guideline limit (Table 1).

Iron, however, poses major aesthetic problems as approxi-

mately 46% of the boreholes have iron concentration higher

than the WHO (2004) guideline limit of 0.3 mg/l, which has

sometimes led to borehole rejection. A major physiological

problem is posed by arsenic, as approximately 21% of the

boreholes sampled have arsenic concentration exceeding

the WHO (2004) maximum provisional guideline limits of

10 lg l–1. It is, however, not as high as those that cause

morbidity in other places such as Bangladesh, India and

China (Nickson et al. 1998; Wang and Huang 1994). The

natural process of oxidation of arsenopyrite appears to be

the main source of arsenic concentration in the boreholes

and hand dug wells in the study area. Though mining en-

hances process of oxidation of arsenopyrite, arsenic can

enter the groundwater system irrespective whether mining

has taken place or not, once the rocks are exposed oxidation

takes place (Carlos et al. 1997).

Effect of land use on groundwater

Gold mining does not seem to have significantly affected

groundwater as evidenced in Fig. 10 where only a few of

the sampled wells are apparently influenced by anthropogic

effect. This may be due to the fact that most communities

from where the wells were sampled are remote from the

mines. Since groundwater circulation is mostly localised

due to the crystalline nature of the geological formations,

the effect of acid mine drainage may not have reached

significant number of wells. Mercury, which is used for

amalgamation has also not been detected in concentrations

higher than the WHO (2004) guideline value. Similarly,

nitrate generated as a result of the use of nitrogen based

explosives in gold mining and agricultural practices do not

have effect on the groundwater quality as evidenced in

Table 2 where the median values of nitrate is 0.2 mg l–1 as

against the WHO (2004) guideline value of 50 mg l–1.

Likewise phosphorous, which is another indicator of pol-

lution from agricultural activities is very low with median

value of <0.1 mg l–1. Thus land use apparently have little

effect on groundwater in the Offin Basin.

Conclusions

The analysis of the hydrochemical composition of ground-

water in the Offin Basin has shown that the groundwater is

generally mildly acidic (median value of pH � 5.9) with few

of the boreholes showing strong acidic character (pH < 4).

Major cations are generally low and Na+ is relatively the

dominant cation while HCO3– is by far the dominant and

forms approximately 82% of the anion content. Principally

the groundwater is either Ca–Mg–HCO3 or Na–Mg–Ca–

HCO3. Other minor water types include Na–HCO3, Na–Cl

and Ca–MgSO4 waters. The groundwater is generally

undersaturated with respect to major carbonate species

namely calcite, dolomite and ankerite. Cation exchange and

dissolution of alumino-silicate minerals particularly pla-

gioclase, amphiboles and pyroxenes by mildly acidic

groundwater is responsible for the major ions in the

groundwater.

The minor or trace metal loading of the groundwater is

not as high as originally anticipated of a mining area. Only

arsenic and iron show significant concentrations in the

boreholes. Approximately 19% of the boreholes have ar-

senic concentration exceeding the WHO (2004) maximum

acceptable limit of 10 lg l–1 for drinking water. However,

these concentrations are not as high as those that cause

morbidity in other places like Bangladesh and India. Nearly

46% of boreholes in the Offin Basin have Fe(total) greater

than 0.3 mg l–1(WHO maximum acceptable limit). Unless

iron is removed through the use of aerators or iron-removal

plants, it poses significant aesthetic water quality problems

for groundwater usage for domestic purposes in the Offin

Basin. Sources of iron in the groundwater include oxidation

of pyrite and arsenopyrite in areas with high sulphides,

reductive dissolution of hematite and goethite in the pres-

ence of organic mater and leaching from ferromagnesian

silicates, such as hornblende and pyroxenes. Land-use

practices apparently have insignificant effect on the quality

of water in the wells in the study area.

Acknowledgments This work was supported by CSIR Water Re-

search Institute. The authors are therefore, grateful to the staffs of

CSIR Water Research Institute, who were also involved in the data

collection for this study but have not been mentioned. We also wish to

thank Mr. Harrison Komladze and Rex Sapaah of the CSIR-Water

Research Institute, Ghana for preparing the Location map.

References

Amasa SK (1975) Arsenic pollution at Obuasi goldmine, town and

surrounding countryside. Environ Health Perspect 12:131–135

Environ Geol (2008) 53:1651–1662 1661

123

Appelo CAJ, Postma D (1999) Chemical analysis of groundwater.

Geochemistry, Groundwater and pollution. Fourth corrected

print. AA Balkema, Rotterdam

Barcelona M, Gibb JP, Helfrich JA, Garske EE (1985) Practical guide

for groundwater sampling. Illinois State Water Survey ISWS

Contract Report 374

Bowell RJ (1992) Supergene gold mineralogy at Ashanti, Ghana:

implications for the supergene behaviour of gold. Miner Mag

56:545–560

Bowell RJ (1993) Mineralogy and geochemistry of tropical rain forest

soils: Ashanti, Ghana. Chem Geol 106:345–358

Bowell RJ (1994) Sorption of arsenic by iron oxides and oxyhydrox-

ides in soils. Appl Geochem 9:279–286

von Bromsse U (1989) Acidification trends in Swedish groundwaters.

Report 3576 National Swedish Environmental Protection Board,

Solna, Sweden

de Caritat P (1995) Intensifying groundwater acidification at

Birkenes, southern Norway J Hydrol 170:47–62

Caritat DP, Danilova S, J�ger Ø, Reimann C, Storrø G (1998)

Groundwater composition near the nickel-copper smelting

industry on the Kola Peninsula, central Barents Region (NW

Russia and NE Norway). J Hydrol 208:92–107

Carlos D, Da Rosa JD, Lyon JS, Udall SL, Hocker PM (1997) Golden

dreams, poisoned streams. Mineral Policy Center, Washington

Claasen HC (1982) Guidelines and techniques for obtaining water

samples that accurately represent the water quality for an aquifer.

U.S. Geological Survey Open File Report 82–1024, 49 pp

Dickson KB, Benneh G (1980) A new geography of Ghana.

Longmans Group Limited, London

Freeze AR, Cherry JA (1979) Groundwater. Prentice-Hall, Inc.

Englewood Cliffs

Gale IN, Robins NS (1989) The sampling and monitoring of

groundwater quality. British Geological Survey. Hydrogeology

Report, No 89/37

Garrels RM, Mackenzie FT (1967) Origin of the chemical compo-

sitions of some springs and lakes. In: Gound RF (ed) Equilibrium

concepts in natural water systems. American Chemical Society

Publications, Washington

Hem JD (1970) Study and interpretation of chemical characteristics of

natural water. US Geol Surv Water-Supply Paper 1473:363

Hem JD (1992) Study and interpretation of chemical characteristics of

natural water. United States Geological Survey Water-Supply

Paper, 2254, United States Government Printing Office, Wash-

ington, Krumbein, W. 1943

Henriksen A (1979) A simple approach for identifying and measuring

acidification of freshwaters. Nature 278:542–545

Henriksen A (1980) Acidification of freshwaters—a large scale

titration. In: Proceedings of international conference on ecolog-

ical impact of acid precipitation. Norway 1980 SNSF Project, pp

68–74

Henriksen A, Kirkhusmo LA (1986) Water chemistry of acidified

aquifers in southern Norway. Water Qual Bull 11(1):34–38 and

61

Hounslow AW (1995) Water quality data analysis and interpretation.

Lewis Publishers, Boca Raton

Jankowski J, Acworth RI, Shekarforoush S (1998) Reverse ion-

exchange in deeply weathered porphyritic dacite fractured

aquifer system, Yass, New South Wales, Australia. In: Arehart

GB, Hulston JR (eds) Proceedings of 9th international sympo-

sium on water–rock interaction. Taupo, New Zealand, 30

March–April1998. Balkema, Rotterdam, pp 243–246

Karanth KR (1994) Groundwater assessment development and

management. Tata McGraw-Hill Publishing Company Limited,

New Delhi. Third reprint

Kesse GO (1985) The mineral and rock resources of Ghana. A.A/

Balkema, Rotterdam, Boston

Knutsson G (1994) Trends in acidification of groundwater. In: Kovar

K, Soveri J (eds) Groundwater quality management. Proceedings

of Tallinn conference in Estonia, September 1993. IAHS publ.

No. 220

Langmuir D (1997) Aqueous environmental geochemistry. Prentice

Hall, Upper Saddle River

McLean W, Jankowski J (2000) Groundwater quality and sustain-

ability in an alluvial aquifer, Australia. In: Sililo et al (eds)

Proceedeings of the XXX IAH congress on groundwater: past

achievements and future challenges Cape Town South Africa

26th November–1st December 2000. A. A. Balkema, Rotterdam

Nkotagu H (1996) The groundwater geochemistry in a semi-arid,

fractured crystalline basement area of Dodoma, Tanzania. J

Afric Earth Sci 23(4):593–605

Nickson R, McArthur J, Burgess W, Ahmed KM, Ravenscroft P,

Rahman M (1998) Arsenic poisoning of Bangladesh groundwa-

ter. Nature 395:338

Parkhurst DL, Appelo CAJ (1999) Phreeqc for windows version

1.4.07. A hydrogeochemical transport model. US Geological

Survey Software

Piper AM (1944) A graphic procedure in the Geochemical

interpretation of water analyses. Am Geophys Union Trans

25:914–923

Smedley PL, Edmunds WM, West JM, Gardner SJ, Pelig-Ba KB

(1995) Vulnerability of shallow groundwater quality due to

natural geochemical environment. Health problems related to

groundwater in the Obuasi and Bolgatanga Areas, Ghana. Report

prepared for ODA under the ODA/BGS Technology Develop-

ment and Research Programme, Project 92/5

Smedley PL (1996) Arsenic in rural groundwater in Ghana. J Afr

Earth Sci 22:459–470

Smedley PL Edmunds WM, Pelig-Ba KB (1996) Mobility of arsenic

in groundwater in the Obuasi gold-mining area of Ghana: some

implications for human health. In: Appleton JD, Fuge R, McCall

GJH (eds) Environmental Geochemistry and Health. Geological

Society Special Publication No. 113 pp 163–181, Chapman and

Hall, London

Stumm (1992) Chemistry of solid-Water Interface: processes at the

mineral–water and particle-water interface in natural systems.

John Wiley, New York. ISBN-0-471-57672-7

Tardy Y (1971) Characterization of the principal weathering types by

the geochemistry of waters from some European and African

crystalline massifs. Chem Geol 7:253–271

UNESCO/WHO/UNEP (1996) Water quality assessments. A guide to

the use of biota, sediments and water in environmental

monitoring, 2nd edn (eds) Deborah Chapman. E & FN Spon,

London, New York

Wang L, Huang J (1994) Chronic arsenism from drinking water in

some areas of Xinjiang, China. In: Nriagu JO (ed) Arsenic in the

environment, Part II: Human health and ecosystem effects.

Wiley, New York, pp 159–172

Williams MW, Smith B (1994) Empirical and modelled hydrochem-

istry of acid mine drainage, iron Duke mine, Mazowe, Zimba-

bwe. British Geological Survey Technical Report WC/94/78/

R.25p

Wilson FH., Hawkins DB (1978) Arsenic in streams, stream

sediments and ground water; Fairbanks area, Alaska. Environ

Geol 2:195–202

Wood WW (1981) Guidelines for collection and field analysis of

ground-water samples for selected unstable constituents. US

Geological Survey Teechniques of Water-Resources Investiga-

tions book 1, ch D2. Us Government Printing Office, Washing-

ton

World Health Organisation (WHO) (2004) Guidelines for drinking

water quality. Final task group meeting. WHO Press, World

Health Organization, Geneva

1662 Environ Geol (2008) 53:1651–1662

123