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Environmental Isotope in Hydrogeology

1. General Introduction

1.1 General

In the last 3-4 decades, environmental isotopes have contributed immensely to studies and investigations in hydrogeology, complementing physical and chemical hydrogeology. Many hydrogeological studies use the stable isotopes of water molecule to determine groundwater quality, origin, recharge mechanism and rock-water interaction. A good number of the applications of environmental isotope in hydrogeology have been in the arid and semi-arid areas of the world, where water scarcity is most acute and pose constraints on economic development. Substantial amount of basic data (1, 2, 3) and results of applied field investigations have already been published (4, 5, 6, 7, 8, 9, 10, 11, 12, 13, 14, 15) on applications of different isotopes for hydrological purposes.

Environmental isotopes provide indications of groundwater age and serving as a natural tracer for groundwater provenance. Stable isotopes carbon, boron, nitrogen and sulphur (i.e. ¹³C/¹²C, ¹¹B/¹⁰B, ¹⁵N/¹⁴N and ³⁴S/³²S) can give valuable information about reactions involving these elements and can also serve as pollution tracers. On the other hand, radioactive isotopes of some of these elements decay, providing us with a measure of the circulation time and invariably groundwater renewability. The most common of the radioisotopes, Carbon-14, is used to estimate groundwater residence time. Isotopes of the uranium series (²³⁴U, ²³⁸U, ²²⁶Ra and ²²²Rn) are also useful tracers in isotope hydrogeology but are often not precise enough to establish the age of groundwater due to mineral-water interactions.

1.2 Historical Background

The history of isotopes dates back to the formation of matter by nucleosynthesis during the birth of our solar system more than 5 billion years ago. This process produced most of the stable and unstable (radioactive) isotopes that naturally exist today (16).

The term “isotope” was first used in 1913 by Soddy Frederick to describe nuclides which occupy the same position in the Periodic Table but which differ in their nuclear properties. Thomson (1913) later showed the element neon to be made up of more than one isotope. In 1919 Francis Aston constructed a mass spectrograph capable of use in the discovery of virtually all elements. But it was not until 1925 before the existence of oxygen of mass 17 was first observed (32). Later on, oxygen-18 and 17 were discovered in natural material (18). Urey in 1931 discovered the hydrogen isotope of mass 2 (called deuterium) and estimated its natural concentration (19 a, b).

With this achievement it became obvious that the isotopic composition of oxygen and hydrogen was highly variable. As more precise measurements of isotope ratios continued, it became evident that almost all D₂O and H₂O depleted in deuterium could be prepared by electrolytic decomposition of water. It was much later that the concentration of pure ¹⁸O was achieved in Switzerland using thermal diffusion method (20). The work of Urey in 1946 formed the basis of isotopic fractionation (21).

Significant achievement in the measurement of natural isotope abundances came with the advent of Nier/McKinney mass spectrometer (22, 23) designed for measurement of small differences in isotope abundance. The discovery of isotopes of water molecules stirred up the several earlier investigations on isotope hydrology as reviewed in the work of Rankama (24). Ever since then investigations involving the use of natural isotope abundances of all lighter elements have continued to increase and many of which has been published in several texts, journals and edited specialist volumes (25, 26, 27, 28, 29, 16, 30, 31, 32, 33).

2. Fundamentals of Isotopes in hydrogeology

Isotopes derived its name from the Greek word: “isos”, meaning equal, and “topos”, which means place (referring to the place in the Periodic Table).

Generally, isotopes of an element are atoms or nuclides having the same number of protons in the nucleus and thus the same atomic number but differing in the number of neutrons and hence in their atomic mass. There are 92 naturally occurring elements comprising more than 1000 isotopes. Most of these occur in terrestrial compounds in trace amounts but some are sufficiently abundant to be determined quantitatively through routine analysis. Hydrogen, for example, is known to have 3 isotopes with the following names and symbols:

(1) H- common Hydrogen (1 proton)

(2) D- deuterium, heavy and stable Hydrogen (1proton + 1 neutron)

(3) T- tritium, radioactive Hydrogen (1 proton + 2 neutrons).

These isotopes can also be described by adding the number of particles in the nucleus of each (i.e. proton + neutron) and placing this at the upper left corner of the symbol for the element. In this way the hydrogen isotopes above may be written as ¹H, ²H, and ³H respectively. For details on element, nucleus and valencies see (30).

Basically, there are two classes of isotopes:

(i) The stable isotopes, which do not change with time, in spite of their concentration being affected by other physico-chemical processes (such as evaporation or condensation).

(ii) The unstable isotopes, which decay with time. The product of this decay are said to be radiogenic if they do not themselves decay. An alternative grouping of isotopes exist:

(a) The environmental isotopes, and (b) The artificial isotopes

Environmental isotopes occur naturally and the investigators have no direct control on the variations of their concentrations while artificial isotopes are those whose variations in the environment are created by man.

The most commonly used environmental isotopes in hydrogeology are the stable isotopes deuterium (²H) and oxygen-18 (¹⁸O) as well as the radioisotope molecules tritium (²H), Carbon-14 (¹⁴C) (see Tables 1 and 2).

Table 1: The stable environmental isotopes (16)

Table 2: The environmental radioisotopes (16)

Note: b^- - beta emission, a - alpha emission, ec – electron capture.

Basically, the characteristics of isotopes useful to hydrogeologists are isotopic ‘signature’ and ‘decay’. Once water reaches the water table, it retains its isotopic content as long as isotope exchange with the reservoir at high temperature does not occur (34).

3. The isotopes: ²H, ¹⁸O, ³H, ¹³C and ¹⁴C

3.1 Stable isotopes of water molecule (²H, ¹⁸O)

Basically, water is composed of hydrogen and oxygen, and therefore occurs with different isotopic combinations in its molecules. As discussed in section 2, the isotopes of hydrogen are: ¹H, ²H, ³H while those of oxygen are: ¹⁶O, ¹⁷O, ¹⁸O. Therefore, the possible stable isotopes of water species are:

H₂¹⁶O, HD¹⁶O, D₂¹⁶O, H₂¹⁷O, HD¹⁷O, D₂¹⁷O, H₂¹⁸O, HD¹⁸O, and D₂¹⁸O.

Significant among these in geochemistry (based on the natural abundance of the isotopes) are ¹H₂¹⁶O, HD¹⁶O, H₂¹⁷O, and H₂¹⁸O. The slight variations in their abundance are caused by small differences in reactivity of the isotopes due to mass differences. Usually, H₂¹⁶O is about 10% lighter than H₂¹⁸O and therefore more reactive. Under closed conditions and without further reactions, the concentrations of isotopes of water molecule remain stable with time.

Generally, stable environmental isotopes are measured as ratios of the two most abundant isotopes of a given element. The absolute abundance ratio of isotopes is not usually measured in natural waters because it requires sophisticated mass spectrometric technique. Due to the low differences, the ratio of the stable isotopes R is given in a delta-notation (in permil units, i.e. parts per thousand) as a deviation from a standard. The mean isotopic composition of seawater, generally known as “SMOW”, (Standard Mean Ocean Water), is used for reference (35). However, the reference now commonly adopted for oxygen and hydrogen stable isotopic variation in natural water is V-SMOW (Vienna Standard Mean Ocean Water), which is isotopically identical to SMOW (36, 9). This difference between samples and the reference standard is expressed in the following relation:

Delta (d)       =             (R_sample - R_standard) / R_standard  x  1000       (permil or  ‰)……….(1)

            (R = ratio of the heavy isotope to the light one, e.g. [¹H²H¹⁶O]/[¹H₂¹⁶O])

By definition the seawater standard has d²H- and d¹⁸O-values equal 0 ‰. Negative values characterize water isotopically depleted ("lighter"), while positive values correspond to water samples isotopically enriched ("heavier") with respect to the standard. The measuring accuracy is 0.15 ‰ for d¹⁸O and 1 ‰ for delta d²H (37). For details of the measuring technique reference is made to IAEA (9, 29).

3.1.1 Isotopic fractionation.

Environmental isotopes of the same element can be partitioned or separated in a thermodynamic reaction due to differences in rates of reaction of the different molecular species. Fractionation is a fundamental process common to stable isotopes of H, B, C, O, N, S and Cl and can occur under equilibrium or non-equilibrium (kinetic) conditions. Fractionation can also occur as a result of molecular diffusion. The different isotopic water molecules have various vapour pressures and freezing points. The changes of the isotope ratio occurring during evaporation, condensation, chemical and biological processes - caused by these differences - are termed "isotopic fractionation".

Vapour derived from seawater is isotopically depleted as compared to SMOW. One of the reasons for the depletion is the lower vapour pressure of the heavy water (e.g. ²H₂¹⁸O). The fractionation or separation occurs at the transition between gaseous, liquid or solid phases, and this is usually expressed by fractionation factor, a, which is defined as:

a = R_reactant  / R_product                                       (2)

For example, the exchange between isotopes of water molecule and the associated fractionation is define as follows:

H₂O_liquid Û H₂O_vapour                                       (3)

a_{liquid - vapour} = (²H/¹H)_liquid  / (²H/¹H)_vapour                                          (4)

or a_{liquid - vapour} = (¹⁸O/¹⁶O)_liquid  / (¹⁸O/¹⁶O)_vapour

Details of theory and applications of stable isotope fractionation exist in publications (34, 38, 16, 39, 30, 40).

Of major importance in isotope fractionation is “kinetic” separation, which occurs when there is a deficit of moisture in the vapour phase. The kinetic separation results from the differing diffusion constants for the heavy (e.g. ²H₂¹⁸O) and the light (e.g. ¹H₂¹⁶O) water molecules: during evaporation the light molecules diffuse faster than the heavy ones through the boundary layer between the water body and the atmosphere (41, 42).

Isotopic fractionation is more efficient if the produced vapour in the process of condensation is constantly removed. This leads to the concept of “Rayleigh fractionation”, a process in which ¹⁸O or ²H is being selectively removed from the vapour phase in such a way that rain becomes progressively lighter in d ¹⁸O and d ²H as it falls farther from the ocean. For details on Rayleigh processes in hydrogeological applications reference is made to (16). Isotope fractionation can be used to interpret isotope data from natural setting and are also useful tool to elucidate and quantify processes as well as reactions in the hydrogeological system.

3.1.2 Dependencies of the isotopic composition.

In environmental isotope studies ¹⁸O and ²H concentrations in precipitation provide a characteristic input signal that varies regionally and over time. The isotopic signatures as encountered in precipitation depend on parameters like temperature, deficit of moisture in the air and the isotope ratio of the water vapour source. When these parameters are taken into consideration, the isotopic signatures give information about the origin of vapour, precipitation, and groundwater - and partly about the climatic conditions during recharge processes in the past. Based on these interrelations some generally valid dependencies of the isotopic composition of a groundwater sample could be deduced (41):

3.1.2.1 The elevation effect:

During the rise of humid air due to orographic obstacles and successive precipitation the concentration of heavy isotopes in the precipitation decreases with elevation. The depletion also is a result of the general decrease of the cloud temperatures with elevation. For the elevation effect the depletion is -1 to -4 ‰ and -0.15 to -0.5 ‰ for d ²H and d ¹⁸O per 100m rise. Elevation correction (also known as altitude or alpine effect) distinguishes groundwater recharged at high altitudes from those of low altitude. It therefore turns out to be a useful tool in hydrogeological studies.

3.1.2.2 The continental effect:

During the condensation of atmospheric vapour the liquid phase (i.e. rain droplets) gets isotopically enriched, the vapour phase gets isotopically depleted. However, the amount of vapour in a cloud is limited. As this process continues, the isotopic signature of the vapour, and consequently of the condensed water, is continuously changing. This, invariably, leads to a situation in which both the precipitation and groundwater are found depleted with respect to heavy isotopes as the distance away from the coast increases.

3.1.2.3 The effect of precipitation rate:

This is otherwise known as amount effect which show the dependence of the isotopic composition on the amount of rainfall: heavier rain effects or greater precipitation amounts result in more negative d ²H and d ¹⁸O values. As the amount of precipitation increases depletion in the rain can be observed. During a single precipitation event significant difference of the isotopic signature are found due to (i) progressive condensation and (ii) variations of the intensity of rain.

3.1.2.4 The temperature effect:

The isotopic composition of precipitation depends on the temperature at which the oceanic water is evaporated into the air. Seasonal fluctuations of the isotopic composition in local precipitation are influenced by fluctuations of temperature. Rain during winter is isotopically lighter than rain during the summer. When in the past the climate was significantly different from those of today then the isotopic signature of groundwater formed in the past will strongly differ from the isotopic signature of modern precipitation and modern groundwater.

Moreover, the dependence of isotopic fractionation on temperature and moisture causes an annual fluctuation (seasonal effect) and depletion with latitude (latitude effect). In arid and semi-arid zones with low moisture saturation and precipitation amounts an enrichment of the heavy isotopes in raindrops occurs while they are falling (evaporation effect). Also due to evaporation surface water like river or lake gets isotopically enriched.

3.1.3 Meteoric Water Line.

In precipitation, rivers, and lakes measured worldwide (35) showed that the delta-values of the stable isotopes fit along a straight line on a d ²H - d ¹⁸O plot. This line, termed the "Global Meteoric Water Line" is characterised by the relation:

            d ²H             =             8d ¹⁸O + 10  ( ‰)                               (6)

Craig’s global line was later refined from more than a decade world-wide monitoring of the stable isotopic composition of precipitation (IAEA Global Network of Isotopes in Precipitation – GNIP, reported in (43)) to be:

d ²H             =             8.13d ¹⁸O + 10.8  ( ‰)                               (7)

From equation (6), the gradient (s) of the GMWL line is 8 and the intercept on the y-axis, “d”, is 10 ‰. The value of d was first used by Dansgaard (34) to characterise the deuterium excess in global precipitation and is defined from equation (6) as:

d = d ²H - 8 d ¹⁸O       ( ‰)                                                        (8)

Changes in gradient of the straight line to values <8 are essentially caused by evaporation during precipitation (see Table 3).

Table 3: Examples of regional meteoric lines (30)

3.2 Tritium (³H)

Tritium is the radioactive isotope of hydrogen, which is produced in a natural manner in the upper atmospheric strata by the influence of cosmic radiation on nitrogen atoms.

¹⁴N + n Þ ¹⁵N Þ¹²C + ³H                                                (9)

Tritium, which in the atmosphere combines with oxygen to form water, may precipitate on earth as rain and thus reach the groundwater. Tritium is usually symbolised as T or simply ³H. Its concentration in water is expressed in Tritium Units (TU): 1 TU corresponds to 1 atom ³H per 10¹⁸ atoms ¹H. Details of measuring techniques and procedure are fully discussed in (44).

Tritium decays with a half-life of about 12.35 years to form ³He. The atmospheric concentration of tritium prior to 1953 was about 3-5 TU. Due to nuclear weapon tests the concentrations in precipitation up to 1963 reached several thousand TU. Hence, since the early 1960s this anthropogenic tritium from bombs was used as tracer to study young groundwater. In the meanwhile, however, in Europe the concentrations have decreased to values <10 TU (45). Measurable tritium in groundwater usually signifies modern recharge. High tritium (>30 TU) indicates recharge in the 1960s while low values (<1 TU) usually signify paleogroundwater (older groundwater) that has mixed with shallow modern groundwater.

Although, qualitative and quantitative approaches to dating groundwater is also possible with tritium (16), the direct age determination of groundwater accurate to the year is somewhat uncertain partly due to the unknown extent of mixing of each year’s recharge with that of the previous year and partly because of high local and temporal variability of the input values in precipitation. However, by measuring ³H together with its daughter ³He, true age determination is possible by calculations not based on the complicated tritium input function.

3.3 Carbon-14 and Carbon-13

There are three isotopes of carbon in nature: common and stable carbon (¹²C), rare and stable carbon (¹³C), and very rare and radioactive carbon (¹⁴C). ¹⁴C is formed, like tritium, in the upper atmosphere from the impact of neutrons produced by cosmic radiation on nitrogen atoms

¹⁴N + n- Þ ¹⁴C + p                                                                (10)

Where n = neutron, p = proton.

The half-life (T_1/2) of carbon-14 is 5,730 ± 30 years. The natural ¹⁴C-level in the atmosphere corresponds to a ratio ¹⁴C/¹²C of 1.18 x 10^-12. The ¹⁴C concentration of a sample is given in pmc units (percent modern carbon) as share of the atmosphere value in 1950, which was fixed to equal 100 pmc. ¹⁴C -values >100 pmc in the atmosphere and in shallow aquifers document the nuclear weapon tests during the 1950s.

The stable isotope ¹³C is often used to determine the initial contents of ¹⁴C. The processes of fractionation can reflect themselves in the ratio ¹³C/¹²C. The¹³C concentration of a sample is given as deviation in  ‰ from a standard. As standard, the ¹³C -value of a marine limestone (Peedee Belemnite = PDB) with the delta-value 0  ‰ is used. Within the atmosphere, d-values range from -7 to -8 ‰. During biological processes the isotopic fractionation is stronger. Hence, biologic CO₂ has values of about -25 -15 ‰ depending on the predominant photosynthesis cycle.

3.3.1 Carbon-dating: principles and problems.

There are two methods of sampling for ¹⁴C: First is by precipitation of approximately 60 millimoles of total carbon as barium or strontium carbonate at a pH >9. The second method involves acidification of water sample, gas stripping the CO₂ with nitrogen, and trapping the evolved CO₂ in a solution of CO₃^-2-free NaOH. Measurements of ¹⁴C are made by beta counting either in gas or liquid phase or by a more recent method using accelerator mass spectrometry of a graphite target.

Radiocarbon activities expressed as percent “modern” carbon (pmc) represent the activity of carbon prior to the dilution by post- industrial ‘dead’ fossil fuel carbonate. For example, a carbon sample having 0 pmc is deem to be dead or have age beyond the limit detectable by radiocarbon dating. The age of a given water sample can be calculated from the relation:

T = 1/l ln C/C₀                                                        (11)

Where t = age (in years), l = decay constant of ¹⁴C, C = measured ¹⁴C activity, C₀ = initial ¹⁴C activity.

The age dating of groundwater with the dissolved, inorganic radioactive carbon (¹⁴C) may be used for ages up to 60,000 years, although poor preservation and subsequent contamination of old material now makes effective dating range shorter. For organic material, effective range is < 50,000 years; for groundwater the range is limited to 30,000 years or less. The main problems associated with ¹⁴C dating are: (i) the determination of the initial C-14 value, (ii) mixture of waters of different age, (iii) diffusive admixture of CO₂ from the atmosphere, (iv) dissolution of carbonates. Despite these problems ¹⁴C -age dating is widespread and yields good results in simple situations (46, 16).

3.3.2 Correction of the carbon-14 age.

By various interactions of recent and fossil carbon an initial ¹⁴C value between 50 and 100 pmc results. Other, partly bacterial processes can also influence the isotopic composition of a groundwater. During the development of the ¹⁴C method several correction models were set up, which reduce the initial activity of samples below 100 pmc. The result is a corrected age. Several correction models exist (46, 47, 48, 49, 50, 51). The resulting ages are often compared to conventional ages, which are obtained with initial values of 100 pmc.

4. Other naturally occurring isotopes

4.1 Chlorine-36 (³⁶Cl)

Chlorine-36 is a radioactive isotope of chlorine whose application to hydrology has attracted much interest in the last decade. It is naturally produced by cosmic rays interacting with the atmospheric argon (⁴⁰Ar), and finds its way into the hydrological cycle either as dry fallout or in precipitation. Thermonuclear bomb testing of 1960s contributed significant amount of ³⁶Cl, thus elevating its concentrations above the natural atmospheric abundance. ³⁶Cl behaves conservatively in most hydrological environment, and like bomb tritium, it is useful in delineating recharge rates. However, unlike tritium, its use for dating modern groundwater is unrealistic. But with a half-life of 3001,000 years, ³⁶Cl is a useful tool in groundwater age determination in the range of 10⁵-10⁶ years.

4.2 Chlorofluorocarbons (CFCs)

Chlorofluorocarbons exist in the atmosphere as: CCl₃F and CCl₂F₂ (simply referred to as CFC-11 and CFC-12 respectively). The production of freons for use as solvents, refrigerants and propellants has released large quantities of these compounds into the atmosphere. The source of CFCs is mainly anthropogenic and studies have shown that their concentration in the atmosphere is steadily increasing since the production of freons started in the mid 1940s. Hydrogeologists have found of a use for CFCs, which would have accumulated as contaminants in the atmosphere, as tracers like tritium. The analysis of CFC compounds is less complicated than that for tritium. This advantage, coupled with the decreasing concentration of bomb tritium (since the 1990s) is responsible for its increasing applications in hydrogeological studies. CFC-11 and CFC-12, with atmospheric residence times of 60 and 120 years respectively, are not isotopes in themselves but equilibrate with water to form a dating tool for groundwater (<50 years old). They are equally applicable as tracers of groundwater and of sewage contamination in water (16, 33).

4.3 Uranium series isotopes

The radioactive decay of uranium and thorium results in the formation of a series of isotopes, which are in themselves radiogenic. Significant among these are: ²³⁴U, ²³⁸U, ²²⁶Ra, ²²²Rn. The uranium series display enormous array of half-lives (10⁵-10⁹) and many geochemical distinctive characteristics that make them useful in hydrogeological applications (16, 33):

(i)                   Evaluation of mixing between groundwater bodies can be achieved on a plot of excess-²³⁴U versus ²³⁸U concentration.

(ii)                 The wide range of half-lives is useful in tracing groundwater movement and investigating geochemical processes.

(iii)