ORIGINAL ARTICLE

Sources of deep groundwater salinity in the southwestern zoneof Bangladesh

Mirza A. T. M. Tanvir Rahman •

Ratan Kumar Majumder • Syed Hafizur Rahman •

Md. Abdul Halim

Received: 29 December 2009 / Accepted: 12 August 2010

� Springer-Verlag 2010

Abstract Twenty groundwater samples were collected

from two different areas in Satkhira Sadar Upazila to

identify the source of salinity in deep groundwater aquifer.

Most of the analyzed groundwater is of Na–Cl–HCO3 type

water. The trends of anion and cation are Cl- [HCO3

- [ NO3- [ SO4

2- and Na? [ Ca2? [ Mg2? [K?, respectively. Groundwater chemistry in the study area

is mainly governed by rock dissolution and ion exchange.

The dissolved minerals in groundwater mainly come from

silicate weathering. The salinity of groundwater samples

varies from *1 to *5%, and its source is possibly the

paleo-brackish water which may be entrapped during past

geologic periods.

Keywords Salinity � Deep groundwater �Rock dissolution � Paleo-brackish

Introduction

In Bangladesh, natural high arsenic groundwater occurs in

shallow aquifers. The problem of the existence natural high

arsenic groundwater in the shallow aquifer has become

progressively more acute with increased dependency on

groundwater for domestic water supplies (Bhattacharya

et al. 1997; Dhar et al. 1997; Nickson et al. 2000). It is

found that 1.2 million wells out of 3–4 million wells are

installed in the shallow aquifers with naturally occurring

high arsenic content since 1960s (DPHE 2001). High

arsenic ([50 lg/L) tube wells are mainly concentrated in

the south-central parts of the country. Out of 64 districts in

Bangladesh, parts of 61 districts shallow aquifers are

occupied with natural high arsenic groundwater and nearly

20–50 million people are at a potential health risk due to

high As in groundwater, and several thousands have

already been affected by chronic arsenicosis (Bhattacharya

et al. 2007; Zahid et al. 2008).

To avoid As problem, peoples in Bangladesh are largely

depend on deep groundwater both for drinking and irriga-

tion purposes; but the scenario at the coastal area is dif-

ferent. Coastal aquifers usually serve as major sources for

freshwater supply, such as drinking or irrigation water. The

deep groundwater in the coastal area is relatively vulner-

able to contamination by seawater intrusion, which makes

groundwater unsuitable for drinking or irrigation (Kim

et al. 2006). In addition, heavy pumping and excessive use

of groundwater near the coast may increase the intrusion of

seawater into the aquifer. The main objectives of this study

are to find out the causes of high salinity in deep ground-

water aquifers of Satkhira area and to identify suitable deep

groundwater aquifers for drinking water supply as well as

for irrigation water.

Location and hydrological setting

The study area is confined to the deep aquifer system of

Satkhira Sadar Upazila, Satkhira District in Bangladesh

(Fig. 1). This area is one of the most foremost places where

the salinization in aquifer would affect the socio-economic

conditions, as well as the environment among all other

places in the coastal area of Bangladesh.

M. A. T. M. T. Rahman � S. H. Rahman

Department of Environmental Sciences,

Jahangirnagar University, Savar, Dhaka, Bangladesh

R. K. Majumder (&) � Md. A. Halim

Institute of Nuclear Science and Technology,

AERE, Savar, GPO Box 3787, Dhaka 1000, Bangladesh

e-mail: ratankm@yahoo.com

123

Environ Earth Sci

DOI 10.1007/s12665-010-0707-z

Three distinct aquifers occur all over the study area

labeled as Aquifers 1, 2 and 3, respectively. Aquifer 1 is

the shallow aquifer overlain by a clay aquitard. Aquifer 2 is

the first deep aquifer separated from Aquifer 1 by a thick

clay aquitard. The aquitard separating the shallow aquifer

from the deep aquifer is a continuous one and it increase in

thickness towards the south. The upper aquitard is very

thick and reaches up to more than 100 m at places. Aquifer

3 is the deepest aquifer separated by deep aquitard from the

upper deep aquifer, i.e., Aquifer 2. All the aquifer systems

are shown in a generalized hydrogeological cross-section

of the study area (DPHE 2006) (Fig. 2).

Sampling and water analysis

A total of 20 deep groundwater samples were collected from

different locations of the study area during August 2008

(Fig. 1). Most of the samples were fitted with a standard

Bangladesh number-6 hand pump. Before sampling, each

well was pumped for at least few minutes until it purged out

approximately twice the well volume or until steady-state

chemical conditions (pH, electrical conductivity and tem-

perature) are obtained. Samples for major analysis were

collected in 500-mL polyethylene bottles. Each bottle was

rinsed with distilled water before pouring the sample water.

The bottles were labeled and air-tight. Two sets of samples

were collected from each location. One was acidified with

0.01 M HCl. The bottles were kept in an opaque bag. The

geographical location of each well was determined with a

handheld global positioning system (GPS) (Explorist 200,

Megellan) and the approximate depths of wells were noted

from the well owner’s records.

The temperature, pH, redox potential (Eh) and salinity

of groundwater samples were measured by EcoScan Ion 6

portable meter, and electrical conductivity was determined

by Hanna HI 8633 portable meter. Total dissolve solids

(TDS), total hardness (as CaCO3), saturation index (SI)

and water type were calculated with Aquachem software

(version 3.7.42).

Concentration of bicarbonate in groundwater sample

was determined with inflection point titration (IPT) method

(US Geological Survey 2006). 50 mL of a sample in a

100-mL beaker was selected for titration. Samples were

titrated by a Digital Titrator (HACH, USA) with incre-

mental additions of 1.6 N H2SO4. When titrating, sample

was stirred using a magnetic stirrer to establish a uniform

mixture of sample and titrant and equilibrium between

sensors and sample. An EcoScan Ion 6 portable meter was

used to measure the pH after the incremental additions of

H2SO4. Titration was done rapidly at first, adding relatively

large acid increments to bring pH to about 5.5; then it was

titrated slowly in small increments until the pH reaches the

value of 4.5. Bicarbonate (mg/L as HCO3-) = 1,000/

mLs 9 (mLa from ip near pH 8.3 to ip near pH 4.5 – mLa

to ip near pH 8.3) 9 CF, where mLsample or mLs, volume

of sample in mm; mLa, volume of H2SO4 added in mm; ip,

inflection point; CF, correction factor for H2SO4 normality

(CF is determined by dividing the actual normality of the

H2SO4 by 0.01639); N, normality of acid solution used.

The authors used calculated Br (Hounslow 1995)

because Br was not measured during the sample analysis.

The ‘calculated Br’, i.e., bromide (Br-) = 0.1634 9

potassium (K?) in mg/L was used as a proxy for Br-. The

use of ‘calculated Br-’ was validated by showing that, for a

subset of wells with complete chemical data, the measured

Br- values (Majumder 2008) plotted against K? gave the

same general results, as when plotted against ‘calculated

Br-’.

Sodium (Na?), potassium (K?), magnesium (Mg2?),

calcium (Ca2?), chlorine (Cl-), sulfate (SO42-) and nitrate

(NO3-) were measured by a Metrohm 761 Compact Ion

Chromatography following standard methods in Kumamoto

Fig. 1 Groundwater sampling site

Fig. 2 Hydrogeological cross-section of the study area along the

lines A–B. Cross-sectional lines A–B is shown in Fig. 1 (modified

after DPHE 2006)

Environ Earth Sci

123

University, Japan. Major-ion compositions are used to

identify the water types with the help of AquaChem software

(version 3.7). The results of physicochemical analyses and

major ionic concentrations of 20 deep groundwater samples

are shown in Table 1.

Results and discussion

General physicochemical characteristics

The salient features of the physicochemical characters and

the correlation matrix of eight chemical parameters of

groundwater samples are presented in Tables 2 and 3,

respectively. The depths of samples vary from 152.4 to

304.8 m, which includes both Aquifers 2 (first deep aquifer)

and 3 (second deep aquifer). These two aquifers are sepa-

rated by an aquitard, which is very thick and reaches up to

more than 100 m at some places. The pH of water is slightly

alkaline that has an average value of 7.61. With an average

electrical conductivity of 1,617.75 lS/cm, it ranges from 397

to 2,650 lS/cm. Salinity of groundwater vary from 0 to 6.4

ppt. TDS has an average value of 861.34 mg/L that range

from 137.3 to 861.34 mg/L. Total hardness shows a wide

range from 90.8 to 640.7 mg/L as CaCO3. All values of the

redox potential (Eh) are negative. Concentrations of Cl- and

Na? show extremely wide range from 5.74 to 956.66 and

15.95 to 621.29 mg/L, respectively. The average value of

HCO3- is 224.32 mg/L that ranges from 147.91 to

320.46 mg/L. The concentrations of Ca2? (25.36–186.53

mg/L) and Mg2? (1.62–67.26 mg/L) are not much high.

SO42- (2.83–24.22 mg/L), NO3

- (3.03–22.42 mg/L) and

K? (1.13–24.22 mg/L) concentrations are very low as

compared to other parameters. It has been seen that

the trends of anion and cation are Cl- [ HCO3- [

NO3- [ SO4

2- and Na? [ Ca2? [ Mg2? [ K?, respec-

tively. It is found that the ratios of Mg2?/Br- and Mg2?/K?

remain to be constant until a magnesium salt begins to pre-

cipitate and that the ratio of K?/Br- remains constant until

potash salts begin to precipitate (Hounslow 1995). Regard-

ing the above-mentioned statement, it is found that 13

samples show Mg2?/K? ratios more or less 1, \ 3 ratios for

6 samples and only one sample show high ratio (results not

shown). It is obvious that in the study area, the precipitation

of ion phases (Mg2? and K?) has a drastic effect on the ion

ratios involving those elements, so that it is reasonable to use

‘calculated Br-’ values in the present study.

The physicochemical parameters of all southern site

samples show low values as compared to the north region.

The spatial distribution of electrical conductivity (Fig. 3)

shows the low values at the southern region than the

northern site of the study area. Meanwhile, the spatial

distribution of Cl- concentrations is similar to that of

electrical conductivity with southward decreasing phe-

nomena. From the correlation matrix (Table 3), it has been

seen that the Na? is strongly correlated with Cl-. Ca2? and

Mg2? show negative correlation with HCO3-.

The Piper (1944) diagram illustrates that 60% of the

sample locates in the marine water region (Fig. 4). This

water is influenced by the saline water. 15% of the sample

is hydrothermal or fossil water. The aquifer of this water

has no source to recharge. The rest of 15% sample water

(S-17, S-18 and S-20) is in the groundwater region which

has rechargeable aquifer by the surface or rain water. These

water samples are located at the southern part of the study

area. Except of these three samples, all other samples are

located at the northern and north-central part of the area.

Although 35% of water is Na–Cl–HCO3 type water, other

65% showed heterogeneous types of water (Table 1).

All Na–Cl–HCO3 water types are located at the northern

part of the study area.

Mechanisms controlling groundwater chemistry

Gibbs’s diagram is widely employed to assess the func-

tional sources of dissolved chemical constituents, such as

precipitation dominance, rock dominance and evaporation

dominance (Gibbs 1970). This diagram represents the ratio

of Cl-: (Cl-?HCO3-) as a function of TDS (Fig. 5). The

distribution of sample data points suggests that the chem-

ical weathering of rock-forming minerals influences the

groundwater quality. Relatively, low saline water (e.g.,

S-17, S-18 and S-20) is more influenced by the rock dis-

solution and most of the samples (except S-3, S-6 and

S-10) are influenced by rock dissolution.

Another important mechanism occurring in groundwater

is ion exchange. The evidence of ion exchange is provided

by the Ca2??Mg2?–HCO3-–SO4

2- versus Na?–Cl-

(meq/L) diagram. Water undergoing ion exchange would

plot along a line with a gradient of -1 on Ca2??Mg2?–

HCO3-–SO4

2- versus Na?–Cl- diagram (Jankowski et al.

1998). The samples plot along the line with a slope of

-0.91, which is approximately equal to one indicating that

ion exchange has taken place in deep groundwater

(Fig. 6a).

Source rock deduction

The cations and anions are mainly derived from rock

weathering rather than evaporation, crystallization and

precipitation according to Fig. 6a. A major proportion of

these ions can be derived from the weathering of crystalline

dolomitic limestones and calcium–magnesium silicates,

chiefly from calcite, gypsum and plagioclase feldspars

(Wen et al. 2005). Calcite (CaCO3) and dolomite

(CaMg(CO3)2) have difficulty in forming when TDS values

Environ Earth Sci

123

Ta

ble

1P

hy

sico

chem

ical

par

amet

ers

of

the

coll

ecte

dg

rou

nd

wat

ersa

mp

les

fro

mth

est

ud

yar

ea

Sam

ple

ID

Dep

th

(m)

Tem

per

atu

re

(�C

)

pH

EC

(lS

/cm

)

Sal

in-i

ty

(pp

t)

TD

S

(mg

/L)

To

tal

har

dn

ess

(mg

/LC

aCO

3)

Eh

(mV

)

Cl-

(mg

/L)

SO

42-

(mg

/L)

NO

3-

(mg

/L)

HC

O3

(mg

/L)

Br-

(mg

/L)

Na?

(mg

/L)

K?

(mg

/L)

Mg

2?

(mg

/L)

Ca2

?

(mg

/L)

Wat

er

typ

e

S-1

21

3.9

72

7.8

7.8

52

,38

04

.38

95

.61

61

.8-

73

.94

01

.27

17

.47

16

.75

30

1.9

73

.39

30

2.0

92

0.7

32

1.2

12

9.8

8N

a–C

l–H

CO

3

S-2

23

7.7

42

7.5

7.8

71

,73

64

.08

01

13

5.3

-7

5.4

35

6.0

51

4.3

31

3.0

43

20

.46

2.7

22

78

.21

16

.66

5.6

44

4.9

2N

a–C

l–H

CO

3

S-3

23

7.7

42

8.1

7.7

92

,40

05

.41

,51

6.7

16

6.0

-7

1.1

94

6.5

92

2.7

81

8.3

22

95

.81

3.8

03

59

.69

23

.26

8.9

25

1.8

2N

a–C

l

S-4

20

92

8.9

7.4

82

,14

04

.98

67

17

6.8

-5

2.4

38

4.1

52

0.4

91

6.5

13

08

.14

3.3

32

71

.37

20

.37

15

.42

45

.43

Na–

Cl–

CO

3

S-5

23

7.7

42

7.6

7.7

61

,59

93

.99

01

.11

77

.7-

68

.54

18

.60

4.3

51

1.0

32

40

.35

2.1

62

85

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13

.20

9.9

05

4.9

0N

a–C

l–H

CO

3

S-6

20

92

8.0

7.7

42

,05

04

.91

,99

8.8

31

2.9

-6

7.7

94

0.9

82

4.2

21

3.5

72

65

.00

3.5

06

21

.29

21

.41

20

.10

92

.26

Na–

Cl

S-7

20

6.0

42

7.3

7.7

41

,85

04

.58

45

.31

15

.0-

67

.64

24

.72

12

.85

12

.31

21

5.7

03

.12

27

1.3

21

9.1

03

.26

40

.73

Na–

Cl–

HC

O3

S-8

23

7.7

42

7.8

7.7

61

,93

44

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47

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07

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69

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24

.40

19

.10

13

.94

25

2.6

72

.51

28

0.2

91

5.3

81

.62

40

.40

Na–

Cl–

HC

O3

S-9

23

7.7

42

7.6

7.7

52

,65

06

.41

,21

2.2

14

6.6

-6

8.6

60

6.7

07

.83

17

.75

30

8.1

42

.71

43

2.2

21

6.5

71

6.5

73

1.4

5N

a–C

l–H

CO

3

S-1

02

24

.03

28

.47

.44

2,0

70

4.5

1,7

74

.36

40

.7-

50

.39

56

.66

20

.80

14

.17

15

4.0

72

.55

38

3.9

11

5.5

84

2.6

11

86

.53

Na–

Ca–

Cl

S-1

12

24

.03

29

.47

.49

2,4

20

5.5

1,1

40

19

8.7

-5

3.2

55

6.7

11

5.7

91

7.1

11

78

.72

3.9

62

69

.92

24

.22

3.1

37

4.4

7N

a–C

a–C

l

S-1

23

04

.82

8.8

7.0

91

,91

03

.77

91

.12

74

.2-

29

.33

50

.77

5.6

41

2.8

01

66

.39

2.7

81

30

.17

17

.00

2.2

91

06

.10

Na–

Ca–

Cl–

HC

O3

S-1

32

37

.74

28

.67

.76

1,0

38

1.9

36

4.9

13

2.9

-6

9.6

13

9.1

82

.93

7.0

22

15

.70

1.0

58

2.9

26

.41

7.2

74

1.2

6N

a–C

a–C

l–H

CO

3

S-1

42

37

.74

30

.27

.65

1,3

29

2.8

60

2.1

18

6.3

-6

3.2

26

3.9

73

.73

9.1

11

47

.91

2.1

61

01

.07

13

.23

17

.88

45

.20

Na–

Ca–

Cl–

HC

O3

S-1

52

74

.32

29

.27

.63

1,4

19

1.2

1,1

93

50

4.3

-6

1.2

63

3.2

31

8.8

92

2.4

21

47

.91

0.1

82

10

.94

1.1

36

7.2

69

1.2

3N

a–M

g–

Ca–

Cl

S-1

62

37

.74

28

.97

.26

97

42

.34

72

.32

56

.6-

39

.41

55

.09

2.8

37

.02

17

8.7

20

.57

28

.63

3.4

99

.70

86

.86

Ca–

Cl–

HC

O3

S-1

72

10

.31

28

.07

.78

39

70

.01

37

.39

0.8

-7

.09

6.6

03

.55

3.5

91

97

.21

0.6

93

5.0

44

.20

6.6

82

5.3

6N

a–C

a–H

CO

3

S-1

81

96

.62

7.9

7.5

54

21

0.0

15

5.2

12

7.7

-5

6.4

5.7

43

.98

3.0

32

15

.70

0.5

11

8.2

33

.12

6.5

04

0.4

7C

a–N

a–H

CO

3

S-1

91

52

.42

8.5

7.2

31

,22

32

.65

48

.42

50

.6-

37

.42

11

.81

9.5

68

.47

15

4.0

71

.61

60

.74

9.8

61

0.1

08

3.7

9C

a–N

a–C

l–H

CO

3

S-2

02

05

.74

28

.07

.48

41

50

.01

62

.71

34

.0-

52

.88

.80

4.7

83

.14

22

1.8

60

.52

15

.95

3.1

86

.54

42

.93

Ca–

HC

O3

Environ Earth Sci

123

are \600 mg/L (Zhang et al. 2007), so when TDS [600 mg/L, Ca2? reacts with HCO3

- to produce calcium

carbonate (Salama et al. 1999).

Ca2þ þ 2HCO3� ! CaCO3 þ CO2 þ H2O ð1Þ

However, precipitation of dolomite is favored by the

presence of alkalinity in the following reaction (Zhang

et al. 2007).

Mg2þþ2HCO3�þCaCO3!CaMg (CO3Þ2 + CO2 + H2O

ð2Þ

Except S-13, S-16, S-17, S-18, S-19 and S-20, all ground-

water samples show TDS values more than 600 mg/L.

Therefore, Ca2? and Mg2? might react with HCO3- pre-

cipitating as calcite and dolomite in Eqs. 1 and 2. Because

Ca2? and Mg2? react with HCO3-, there would be straight

positive correlation between Ca2?/HCO3- and Mg2?/

HCO3- ratios. However, there is no such positive correla-

tion (r2 = -0.05) between Ca2?/HCO3- and Mg2?/HCO3

-

Ta

ble

2T

he

sali

ent

feat

ure

so

fth

ep

hy

sico

chem

ical

char

acte

rso

fth

eg

rou

nd

wat

er

Dep

th

(m)

Tem

per

atu

re

(�C

)

pH

Co

nd

uct

ivit

y

(lS

/cm

)

Sal

init

y

(pp

t)

TD

S

(mg

/L)

To

tal

har

dn

ess

(mg

/Las

CaC

O3)

Eh

(mV

)

Cl-

(mg

/L)

SO

42-

(mg

/L)

NO

3-

(mg

/L)

HC

O3-

(mg

/L)

Na?

(mg

/L)

K?

(mg

/L)

Mg

2?

(mg

/L)

Ca2

?

(mg

/L)

Min

15

2.4

27

.37

.09

39

70

13

7.3

90

.8-

7.0

95

.74

2.8

33

.03

14

7.9

11

5.9

51

.13

1.6

22

5.3

6

Max

30

4.8

30

.27

.87

2,6

50

6.4

1,9

98

.86

40

.7-

75

.49

56

.66

24

.22

22

.42

32

0.4

66

21

.29

24

.22

67

.26

18

6.5

3

Av

g2

26

.61

28

.32

7.6

11

,61

7.7

53

.38

86

1.3

42

14

.82

-5

6.7

24

09

.60

11

.79

12

.05

22

4.3

22

21

.96

13

.40

14

.13

62

.80

MD

23

0.8

92

8.0

57

.70

1,7

93

3.9

58

46

.51

76

.8-

62

.23

92

.71

11

.21

12

.92

21

5.7

02

70

.62

15

.48

9.3

14

5.3

1

Min

min

imu

m,

Ma

xm

axim

um

,A

vgav

erag

e,M

Dm

edia

n

Table 3 Correlation matrix of the eight chemical parameters of

groundwater samples

Na? K? Mg2? Ca2? Cl- SO42- NO3

- HCO3-

Na? 1.00

K? 0.73 1.00

Mg21 0.25 -0.21 1.00

Ca2? 0.22 0.05 0.49 1.00

Cl- 0.88 0.64 0.45 0.51 1.00

SO42- 0.75 0.62 0.39 0.30 0.80 1.00

NO3- 0.69 0.60 0.50 0.21 0.77 0.74 1.00

HCO3- 0.50 0.47 -0.27 -0.52 0.19 0.35 0.28 1.00

Fig. 3 Spatial distribution of electrical conductivity of groundwater

Environ Earth Sci

123

ratios. Moreover, Ca2? and Mg2? showed negative corre-

lation with HCO3- (Table 3). Considering this fact, it is

evident that there should not be much carbonate or dolomite

dissolution in the studied groundwater. By examining indi-

vidual effect of carbonate or dolomite dissolution, it is found

to have only negligible positive correlation between Mg2?

and HCO3- (r2 = 0.02) and between Ca2? and HCO3

-

(r2 = 0.23). This result also complies with the above-men-

tioned statement.

Saturation index for calcite and dolomite

The state of saturation of groundwater sample with respect

to important carbonate mineral (calcite and dolomite) is

determined using the calculated data produced by the

Aquachem. The SI of a given mineral is defined as

(Kortatsi 2006):

SI ¼ log10ðIAP=KspÞ ð3Þ

where IAP is the ion activity product and Ksp is the solu-

bility product at a given temperature (i.e., the thermody-

namic equilibrium constant adjusted to the temperature of

the given sample).

Computed saturation indices for calcite and dolomite are

plotted in Fig. 6b. In Fig. 6b, dolomite and calcite satura-

tion indices are indicated on the x-axis and y-axis,

respectively. The four quadrants of the plotting field out-

side the equilibrium area labeled A–D represent different

kinds of non-equilibrium conditions with respect to both

carbonate species. Quadrant A represents super-saturation

with respect to both dolomite and calcite. This super-

saturation condition probably arises when water that is

brought to equilibrium with these carbonate species is

subsequently transported in a different environment where

a higher pH is attained due to the loss of CO2 or other

means (Langmuir 1991). Only five samples, i.e., 25%

samples (S-1, S-4, S-5, S-18 and S-19) are placed in

quadrant A precipitated as calcite, or dolomite. Quadrant B

represents super-saturation with respect to calcite, but

under-saturation with respect to dolomite. Analyses plot-

ting in this field represent water that is undergoing incon-

gruent dissolution of dolomite and/or precipitation of

calcite. Three samples, i.e., 15% samples (S-3, S-8 and

S-13) are plotted in quadrant B, which are from the same

depth.

Quadrant C represents under-saturation with respect to

both calcite and dolomite. An analysis plotting in this

quadrant probably represents water that has come from an

environment where calcite and dolomite are impoverished

or where Ca2? and Mg2? exist in other forms. Groundwater

that has not reached equilibrium with the carbonates

because of short residence time would also probably plot in

this quadrant (Langmuir 1991). Waters of this type will

dissolve calcite if the water comes into contact with a

calcite or dolomite source. Except the samples fallen in

quadrants A and B, all other groundwater samples are

plotted in quadrant C, i.e., 40% of all samples. Quadrant D

represents super-saturation with respect to dolomite and

Fig. 4 Piper plot of analyzed groundwater sampler (arrow sign

indicates increase in salinity)

Fig. 5 Mechanism controlling groundwater quality (after Gibbs

1970)

Environ Earth Sci

123

under-saturation with respect to calcite. None of the water

samples are plotted in this field. From this discussion, it can

be concluded that carbonate and dolomite are not found

much in the sample waters. This means carbonate and

dolomite may not be the dominate source of Ca2?, Mg2?

and HCO3- in the study area.

Silicate weathering

The plot of Ca2??Mg2? versus HCO3-?SO4

2-shows that

most of the data (except S-10 and S-15) locate below the

1:1 trend, reflecting the requirement of cations from

weathering of silicate rocks as shown in Fig. 7a.

The plot of Ca2??Mg2? versus HCO3- in Fig. 7b

shows that most of the data fell below the 1:1 trends,

although some points approach this line. The low

(Ca2??Mg2?)/HCO3- ratios (\0.5 or below 1:2 line) can

be the result of either Ca2??Mg2? depletion by cation

exchange or HCO3- enrichment, i.e., for S-1, S-2, S-3, S-4,

S-5, S-7, S-8, S-9, S-13, S-17, S-18 and S-20. However,

high ratios cannot be attributed to HCO3- depletion, under

the existing alkaline conditions, HCO3- does not form

Fig. 6 Bivariate plot of

a Ca2??Mg2?–HCO3-–SO4

2-

versus Na?–Cl-. b Relation

between calcite and dolomite

saturation indices

Fig. 7 Bivariate plots of

a Ca2??Mg2? versus

HCO3-?SO4

2-, b Ca2??Mg2?

versus HCO3-, c Ca2??Mg2?

versus total cation and

d Na??K? versus total cation

Environ Earth Sci

123

carbonic acid (H2CO3) (Spears 1986). The samples (all

except S-10, S-12, S-15 and S-19) below the 1:1 line

suggest that the excess of alkalinity of these waters bal-

anced by alkalis Na??K?. The excess of alkaline earth

element (Ca2??Mg2?) over HCO3- in some samples (i.e.,

S-10, S-12, S-15 and S-19) reflect an extra source of Ca2?

and Mg2? and is balanced by the Cl- and SO42- (Wen

et al. 2005). Further, the plot of Ca2??Mg2? versus total

cation shows that the data are far below the 1:1 trend,

reflect an increasing contribution of Na? and K? as TDS

increase (Fig. 7c).

The Na? dominance, an index of weathering, suggests

that the ions result from silicate weathering or dissolution

of soil salts (Wen et al. 2005). Groundwater in the area has

a higher ratio (averaged 0.74) of (Na??K?) versus total

cation (Fig. 7d), depicting the contribution of cations via

silicate weathering and soils, to some extent (Sarin et al.

1989). This high ratio also shows that silicate minerals

might come into the solution by the dissolution of plagio-

clase (albite) to kaolinite, montmorillonite and/or gibbsite

according to Eqs. 4, 5 and 6, respectively.

2NaAlSi3O8Albite

þ2CO2 þ 11H2O

¼ Al2Si2O5ðOHÞ4Kaolinite

þ2Naþ þ 2HCO3� þ 4H4SiO4 ð4Þ

2NaAlSi3O8Albite

þMg2þ þ 4H2O

¼ 2Na0:5Al1:5Mg0:5Si4O10ðOHÞ2Montmorillonite

þ2Naþ þ H4SiO4 ð5Þ

NaAlSi3O8

AlbiteþCO2 þ 8H2O

¼ Al(OH)3Gibbsite

þNaþ þ HCO3� þ 3H4SiO4: ð6Þ

The Na-normalized (Gaillardet et al. 1999) ratios for

Ca2? and Mg2? may have related to each other.

Accordingly, in the plot of the molar ratios of Mg2?/Na?

versus Ca2?/Na? are shown in a log–log plot in Fig. 8a.

Recharging waters flowing through carbonates-rich aquifer

show high Ca2?/Na? and Mg2?/Na? rations. The end

member having lower Na-normalized ratios is that of

waters-draining silicates. The molar Ca2?/Na? ratio of

average crustal continental rocks is close to 0.6 (Taylor and

McLennan 1985), and due to the higher solubility of Na?

relative to Ca2?, lower Ca2?/Na? molar ratio are expected

in groundwater, which are related to weathering of

silicates. In Fig. 8a, the observed groundwater with low

Ca2?/Na? molar ratios is being influenced by silicate

weathering rather than carbonate dissolution. Similarly, the

plot for Ca2?/Na? versus HCO3-/Na? molar ratios in

Fig. 8b, low molar ratios of Ca2?/Na? and HCO3-/Na? for

the groundwater samples are the indication of silicate

weathering.

Source of the salinity

The salinity ranges from 0 to 6.4 ppt with an average of

3.38 ppt in the study area. The highest salinity has been

observed at north of S-9 (depth of 237.74 m) and the

lowest is for S-17, S-18 and S-20 all of which are at south.

The groundwater salinity gradient should have an

increasing trend from the north of Bangladesh to the south

near the Bay of Bengal. However, it is observed that deep

groundwater salinity decreases towards south in the study

area, which is similar to the spatial distribution pattern of

electrical conductivity (Fig. 3) and Cl- concentrations.

Most of the high salinity water is found within a depth

ranged from 200 to 250 m.

In the study area, the potential salinization sources in

deep groundwater are diverse, including natural saline

groundwater, halite dissolution and the presence of paleo-

brackish water and seawater. However, there are no signif-

icant accumulations of evaporates in the hills adjoining the

Bengal Basin (Krishnan 1985), so halite dissolution may not

be the possible source of deep groundwater salinity in the

study area. Meanwhile, the TDS is an important parameter

that can be used to observe the influence of major compo-

nents in groundwater quality. The groundwater concentra-

tions of Na? and Cl- are plotted against TDS (Fig. 8c, d).

The plot shows that most Na? and Cl- ions are positively

correlated (r2 = 0.82, 0.97, respectively) with TDS

(Fig. 8c, d). Except the samples S-15, S-16, S-17, S-18, S-19

and S-20, all others are on the marine origin site of Piper

diagram (Fig. 4). In Fig. 8e, Na? \ Cl- indicates reverse

softening, i.e., the presence of brine or seawater (Hounslow

1995).

According to Appelo and Postma (1999), the chemical

reactions during fresh/salt water displacements can be

deduced more specifically by calculating a composition

based on the conservative mixing of salt water and fresh

water, and comparing the conservative concentrations with

those actually found in the water analysis. The concentra-

tion of an ion i, by conservative mixing of seawater and

fresh water is

mi;mix ¼ fsea � mi;sea þ ð1� fseaÞ � mi;fresh ð7Þ

where, mi is the concentration of i (mmol/L), fsea is the

fraction of seawater in the mixed water and subscripts mix,

sea and fresh indicate the conservative mixture and end

member seawater and fresh water.

The fraction of seawater is normally based on the Cl-

concentration of sample. Cl- is assumed to be a conser-

vative parameter to find the seawater contribution to

rainwater. The Cl--based fraction of seawater is

fsea ¼ ðmCl;sample � mCl;freshÞ=ðmCl;sea � mCl;freshÞ ð8Þ

Environ Earth Sci

123

S-20 is of Ca–HCO3 type which is same as the rain

water type of Bangladesh (Majumder 2008). Therefore,

when considering the Cl- concentration of S-20 as the

freshwater, it is compared to get the mixing proportion

between the seawater and fresh water. From this equation,

it has been found that the fractions of seawater are *1%

for S-13, S-14, S-16 and S-19, *2% for S-1, S-2, S-4, S-5,

S-7, S-8 and S-12, *3% for S-9, S-11 and S-15, *5% for

S-3, S-6 and S-10. However, samples S-17, S-18 and S-20

do not have any seawater fractions.

The hydrochemistry of some minor elements, such as

bromine (Br) and strontium (Sr), is a good marker of the

mixing between fresh water and seawater. The Br-/Cl-

ratios have been used to distinguish salinity of marine and

non-marine origins (Andreasen and Fleck 1997). Yongje

Kim et al. (2003) observed that the Br-/Cl- ratios of

groundwater with high Cl- concentrations are maintained

in the normal seawater range (0.0033–0.0037) with the

mean value of 0.0036, indicating that groundwater salini-

zation results mainly from mixing with seawater. In the

present study, only two samples (S-6 and S-16) show

evidence of mixing with seawater having the Br-/Cl-

ratios of 0.0037 (Fig. 8f). However, when seawater

intrudes into fresh coastal aquifer, CaCl2 or MgCl2 type

water results (Appelo and Postma 1999). In this case, Na?

of seawater is being replaced with either Ca2? or Mg2? of

the clay minerals whereby Na? is being adsorbed onto the

clay mineral surface according to Eqs. 9 and 10

Fig. 8 Bivariate plots of

a Mg2?/Na? versus Ca2?/Na?,

b Ca2?/Na? versus

HCO3-/Na?, c Na? versus

TDS, d Cl- versus TDS, e Na?

versus Cl- and f Br-/Cl- versus

Cl-

Environ Earth Sci

123

ðCaþ2 þ Mgþ2Þ � 4ðHCO3�Þ þ 2ðNaXÞ

¼ CaX2 þMg(HCO3�Þ2 þ 2NaHCO3

� ð9Þ

2NaClþMgX2 ¼ 2NaX þMgCl2 ð10Þ

where X signifies the exchanger. Interestingly, no samples

show the CaCl2 or MgCl2 type of water (Table 1). There-

fore, seawater intrusion is not the factor of salinization in

the study area.

The sea level in the Bengal Basin has been changed

from the past. During the holocene period, about 6 Ka is

the peak period of the last highest sea level (DPHE 2006).

The earliest Ganges delta development phase took place

about 5–2.5 Ka (Allison et al. 2003). Majumder (2008)

observed that the age of the deep groundwater fell along

the seawater line ranged from *6 to 25 Ka. As a result,

brackish water originated from the sea is trapped within the

aquifer of the study area, which is once exposed to the sea.

The current salinity definitely originates from this saline

aquifer, because no other possible sources of salinization

are observed in the study area. Sikdar et al. (2001)

observed similar type of brackish connate water pockets in

the western site of Bengal Delta.

Conclusion

The high salinity along with the high electrical conduc-

tivity, sodium and chloride concentrations are observed in

the study area. Rock dissolution and ion exchange are the

main mechanisms that control the groundwater chemistry

in the study area. The dissolved minerals in deep ground-

water mainly come from the silicate weathering. Carbonate

dissolution is almost negligible in the deep aquifer. The

source of salinity in deep groundwater is the paleo-brackish

water, which may be entrapped during past geologic peri-

ods. From this research work, the best aquifer has been

identified at the southern site of the study area which could

be developed as potential aquifer for the sustainable water

usages.

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