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Environmental Geochemistry and Isotopes 1
Part II Tracing the Origin and Geochemical
Evolution of Natural Waters
Chapter 4: Tracing the Water Cycle In groundwater resource and contamination studies, certain questions of its origin and subsurface history come to mind. The locality and timing of recharge is basic to water resources issues such as renewability and susceptibility to contamination. The subsurface flow path, mixing and geochemical evolution of groundwater greatly impact its quality. Physical parameters such as discharge measurements, and water levels of rivers and groundwaters, when interpreted in conjunction with meteorological data, are fundamental to water resource evaluation. However, the stable isotopes of water provide complementary tools for distinguishing different water sources, identifying areas and mechanisms of recharge, and calculating mixing. Before studying the geochemical evolution of groundwater, let’s take a look at tracing the water itself. The natural isotopes in water, 18O and D, are an intrinsic component of the water molecule. They are selectively partitioned at each step of the hydrological cycle, from primary evaporation over the oceans, through condensation and precipitation to groundwater recharge and runoff back to the seas. Isotopes provide characteristic fingerprints of water’s origin and history. This fingerprint is the basis of isotope hydrology. The principal hydrological processes that affect the isotopic fingerprint of water are:
1) primary formation of atmospheric vapor by evaporation under different humidity conditions 2) condensation in clouds and rainout along a temperature gradient, imparting a gradual depletion on
both the vapor and successive precipitations. 3) evaporation from soils and surface waters – which enriches the residual water in both isotopes. 4) mixing of waters in the surface or subsurface – a process that often occurs in open boreholes which
can connect and sample both shallow and deep aquifer zones 5) isotope diffusion through clay aquitards – a form of mixing that alters the isotope signature of pore
waters by diffusive fractionation. 6) rarely . . . isotope exchange during mineral–water and gas–water reactions
The last process of isotope exchange with minerals and gases is important in only very exceptional circumstances of low water/rock or water/gas ratios. While not uncommon, neither isotope exchange or diffusion are effects that would be observed in freshwater resources. Thus, with the exception of deep geological settings (e.g. geothermal waters and brines) or highly contaminated waters (e.g. landfills), it is safe to assume that the isotope content of groundwater has not been affected by subsurface processes beyond simple mixing. Considering the complexity of the hydrological cycle, it may be surprising that 18O and D behave at all in a predictable fashion. Yet two fundamental correlations are observed for fresh waters, which are a product of the first two hydrological processes listed above. These correlations are:
2 Chapter 5 Tracing the Hydrological Cycle
-55
-45
-35
-25
-15
-5
5
-50 -25 0 25
N. Atlantic coastal
Island stations
Greenland
Antarctica
Mean Annual Air Temperature (°C)1
8O
‰ S
MO
W
W
δ18O = 0.69 T – 13.6
δ2H = 8.13 δ18O + 10.8
-200
-160
-120
-80
-40
0
40
-25 -20 -15 -10 -5 0
δ18O ‰ VSMOW
δ2H
‰ V
SM
OW
w
Warm regions
Cold regions
Global Precipitation
• a strong linear correlation between δ18O and δD in meteoric waters on a global and local scale (Fig. 4-1A)
• the strong correlation between δ18O (and δD) values for precipitation and mean annual air
temperature (MAAT) (Fig. 4-1B; modified from Dansgaard, 1964)
Fig. 4-1 A — The δ18O–δD correlation for global precipitation plotted from data on the International Atomic Energy Agency GNIP
database (http://isohis.iaea.org/). Craig (1961) first observed this correlation in global freshwaters. B— The global T–δ18O relationship
for precipitation, modified from Dansgaard (1964). Temperature is mean annual surface air temperature (MAAT) at the station. For
deuterium, Dansgaard found δD = 5.6 Tannual – 100‰ SMOW. Note that SMOW refers to the laboratory reference water used at that
time. It has since run out and been replaced by VSMOW.
THE TEMPERATURE – δδδδ18
O CORRELATION IN PRECIPITATION
In the early days of isotope hydrology research, Harmon Craig, an American pioneer of stable environmental isotope research, observed that fresh waters in warm regions were enriched in 18O and D while those from cold regions were isotopically depleted. His Danish colleague, Willi Dansgaard, had begun measuring the isotopic composition of precipitation over a range of latitudes from northern
Greenland to Antarctica and correlating them with mean annual air temperature. His temperature-δ18O correlation (Fig. 4-1) is fundamental to isotope hydrology. Why? If the isotopic contents of precipitation, and so of surface waters and groundwaters, are controlled by the temperature during precipitation, then we can distinguish waters originating from precipitation during different seasons, different elevations, across a range of latitudes, or even through changing climates. How is this correlation produced? The answer lies in the mechanism of precipitation or rainout, a process driven by the continual cooling of a vapor mass.
Rainout and Rayleigh distillation of 18
O Atmospheric water vapor is generated by primary evaporation of ocean water, with contributions from open water bodies and plant transpiration along the air mass trajectory. Precipitation is generated through cooling of the vapor mass. The temperature of the vapor must drop below the dew point (where humidity = 100%) for condensation to occur. Cooling occurs by adiabatic expansion (no loss of enthalpy) as warm air rises by convection or orographically over mountains or over more dense air masses. It also cools by radiative heat
Environmental Geochemistry and Isotopes 3
loss. Only through cooling will condensation and precipitation occur. This is the process of “rainout”, and it is driven by decreasing temperature. Within the cloud, fractionation between vapor and water or vapor and ice preferentially partitions 18O and D into the rain or snow. In this way, 18O and D are distilled from the vapor phase during rainout. The vapor becomes progressively depleted in 18O and D along the vapor mass trajectory, following a Rayleigh–type distillation as discussed in Chapter 2. As the residual vapor becomes depleted in 18O and D, subsequent rains will also become progressively depleted, although still enriched with respect to their condensing vapor phase. Rainout is an evolution towards colder, isotopically-depleted precipitation (Fig. 4-2).
Fig. 4-2 The change in the 18O content of precipitation according to a Rayleigh distillation.
We can simulate the isotopic evolution of precipitation during rainout using a simplified Rayleigh distillation equation (modified from above in Chapter 2):
δ18Ow(f) = δo
18Ov – ε18Ov–w (1 + ln f ) Rayleigh function in δ–‰ notation
where δo18Ov is the isotope value for the initial water vapor, f is the residual fraction of the vapor reservoir
in the air after a certain amount of rainout, and δ18Ov(f) is the isotopic value for the residual fraction, f. The
enrichment factor ε18Ov–w is for equilibrium water-vapor exchange at the prevailing in-cloud temperature.
At 25°C, ε18Ov–w is –9.3‰ (Table 2-7).
Example 4-1 Isotopic evolution during rainout.
A vapor mass has a δ18O value of –13‰ and a temperature of 25°C as it begins to rain. What is the initial
δ18O composition of the rain at 25°C and after cooling to 25°C and losing 25% of its mass?
δ18O of initial rain (25°C):
δ18Ov(f=1) = –13‰
ε18Ow-v = 9.3‰ @ 25°C (Table 2-7)
δ18Orain(f=1) = δ18Ov(f=1) + ε18Ow-v
4 Chapter 5 Tracing the Hydrological Cycle
= –13 + 9.3 = –3.7‰
δ18Orain after rainout has proceeded by 25% by cooling to 20°C:
ε18Ow-v =– 9.7 at 20°C (equation with Table 2-7)
f = 1 – 0.25 = 0.75
δ18Ow(f) = δo
18Ov – ε18Ov-w (1 + ln f )
δ18Ov(0.75) = –13 – (–9.7 · (l – 0.29))
= –6.1‰
δ18Orain after rainout has proceeded by 50% by cooling to 10°C:
f = 0.5
ε18Ov-w = –10.6 (equation with Table 2-7)
δ18Ow(f) = δo
18Ov – ε18Ov-w (1 + ln f )
δ18Ov(0.75) = –13 – (–10.6 · (l – 0.69))
= –9.79‰
In this example, the δ18O composition of the rain has evolved from –3.7‰ to –9.7‰ during rainout of 50% of the water vapor reservoir (f = 0.5). Note that the decreased temperature actually increases the water-vapor enrichment of 18O, which acts to enrich the rain as temperature decreases. However, the distillation of heavy isotopes from the vapor has a stronger effect, and results in an overall depletion at decreased
residual fractions, f.
It is interesting to plot the isotopic evolution of precipitation during rainout, based on the simple
calculations in Example 4-1. This is done in Fig. 4-3, using the same initial conditions from Example 4-1.
The first plot shows the isotopic evolution of rain and snow which forms from a continually cooling vapor
mass. Note the offset for precipitation when the condensation temperature drops below freezing, due to
difference between water-vapor and ice-vapor fractionation at 0°C.
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0
00.20.40.60.81
-450
-350
-250
-150
-50
f
δ18
O
Snow
Rain
Vapor
δD
-50
-40
-30
-20
-10
0
-20-1001020
T°C
-450
-350
-250
-150
-50Rain
SnowVapor
δ18
O δD
Fig. 4-3 The isotopic evolution of rain and snow during rainout. Plot on left shows the gradual depletion for both 18O and D. Both isotopes follow the same depletion (but on different scales) with decreasing residual vapor fraction f. The plot on right show the ideal
evolution of δ18O and δD for a single vapor mass, showing the co-depletion in 18O and D with decreasing temperature T. Note the
jump to higher δ18O and δD at the rain-to-snow transition due to the greater fractionation between vapor and ice vs. vapor and water.
From these plots, it is clear that rainout is a powerful mechanism for partitioning isotopes. This evolution
towards precipitation with depleted δ18O and δD values occurs along weather trajectories on a continental
scale, but can also occur during a single rainfall event. The isotopic composition of rainfall at the end of a
storm is invariably more depleted than the rain at the outset, and often by a few permil.
Environmental Geochemistry and Isotopes 5
However, weather is a complex phenomenon, and the evolution that presented in Fig. 4-3 is an ideal system only. Most weather systems acquire new sources of vapor along their paths. The precipitation collected at a given meteorological station can also come from different origins, depending on the season and the track of any given storm event. Further, the in-cloud temperature of condensation (which controls fractionation) is seldom the same as the air temperature measured on the ground. For these reasons, the correlation between
MAAT and δ18O (and D) of precipitation for any given site is based on the average of long term monitoring
data. Plotting temperature vs δ18O at the scale of individual precipitation events generally produces a very poor correlation.
A global map of average annual δ18O in precipitation illustrates the partitioning of isotopes between cold
and warm regions was created from mean annual precipitation data collected within the IAEA-World
Meterological Organization survey of precipitation, using a geographical information system (GIS) for
contouring. On this global scale, the partitioning of 18O into warmer, low-latitude precipitation is clear.
Fig. 4-4 Global map of δ18O for precipitation, based on annual averaged data from collected from International Atomic Energy
Agency (IAEA) stations over the past 30 years (Rozanski et al., 1993). These long term data are available at the IAEA website <http://isohis.iaea.org/ >.
THE METEORIC WATER LINE FOR δδδδ18
O AND δδδδD IN METEORIC WATERS
The second key feature of 18O and D in meteoric waters is their strong positive correlation. As discussed above, this was first noted and published by Craig (1961) for global fresh waters. The regression line for these data give is the “global meteoric water line” (GMWL), defined by Craig as:
δD = 8 δ18O + 10 ‰ SMOW The importance of Craig’s findings is that the meteoric water line provides a datum to compare the isotopic composition of surface and groundwater. This allows the isotope hydrologist to derive information such as recharge conditions, evaporative losses and mixing.
Slope of the meteoric water line The GMWL is actually an average of many local or regional meteoric water lines, which differ from the global line due to varying climatic and geographic parameters. Local lines will differ from the global line in
6 Chapter 5 Tracing the Hydrological Cycle
both slope and deuterium intercept (δD value at δ18O = 0) (Fig. 4-5). The correlation between δ18O and δD is due to the similar behavior of both isotopes during rainout. The slope of this regression reflects the greater fractionation for D, which is about 8 times greater than that for 18O. In fact, the ratio of the two enrichment factors is equal to the global slope, s, of 8 at 30°C, and increases to about 9 at 0°C, as shown in Fig. 4-5.
v-w18
v-w
Oε
εD=
3.9
76= 8.17 @ 25°C
= 6.11
106= 9.14 @ 0°C
-300
-200
-100
0
-40 -30 -20 -10 0
δ18O ‰
D ‰
Snow
Rain s = 8.2
s = 9.0
s = 9.5
-160
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-80
-40
0
40
-20 -15 -10 -5 0 5
Global · · · · · · ·
EMWL · · · · · · ·
Victoria · ·
Ottaw a · · · · · ·
Midw ay Is. · · · ·
δ18O ‰ VSMOW
D ‰
VS
MO
W W
δ2H=8 δ18O+10
δ2H=8 δ18O+22
δ2H=7.5 δ18O–2
δ2H=7.6 δ18O+7
δ2H=6.8 δ18O+6
Fig. 4-5 Left: meteoric water line for δ18O and δD in precipitation following the ideal Rayleigh distillation shown in Fig. 4-3. Right:
Regional meteoric water lines for selected stations in the IAEA global network for isotopes in precipitation (GNIP). The Eastern Mediterranean line (EMWL) has a high deuterium excess due to low humidity conditions during primary vapor formation in that region. Victoria, B.C., Canada has a very low deuterium intercept due to high humidity conditions in the western Pacific where vapor
originates. The lower slope of the regional meteoric water lines (s = 6.8 to 8) compared with the ideal rainout curves (s = 8.2 to 9.5) reflects the non-ideal rainout processes of regional precipitation that include gains of vapor along the way and secondary evaporation during and after rainfall.
This slope can be affected by evaporation that occurs after condensation. If rain is falling through a dry air column above the ground, some will evaporate, imparting a kinetic fractionation on the drop. Dansgaard (1964) described such evaporation of rain as the amount effect on the isotopic composition of precipitation. In particularly arid regions, local meteoric water lines calculated using data from all the rain events are characterized by a slope that is closer to and even less than 7. However, based on data from only significant rain events (say >20mm), which can saturate the air column and minimize evaporation of droplets, the slope becomes closer to 8.
Environmental Geochemistry and Isotopes 7
-60
-40
-20
0
20
40
-8 -6 -4 -2 0 2 4 6 8
δ18O ‰ VSMOW
δD
‰ V
SM
OW
W
Northern Oman meteoric w ater line
rains >20mm (δ2H = 7.5 δ18O + 16.1)
All rains
δ2H = 5.2 δ18O + 10
Fig. 4-6 The isotopic composition of precipitation from an arid region (Oman) showing the amount effect where evaporation during rainfall can enrich its isotopic composition.
Deuterium excess and the origin of water vapor
The second parameter defining a meteoric water line is the deuterium intercept. In Fig. 4-5 the deuterium
intercept of the different regional meteoric water lines varies between 22 and less than 0. This is partly due
to the variations in slope, which affects the intercept on the δD axis, but is also affected by a uniform offset
of the line with respect to δD. Dansgaard (1964) defined this off-set as the deuterium excess, by using a
standard slope of 8. Accordingly, the deuterium excess or d of precipitation is calculated for any given 18O -
D isotope pair as:
d = δD – 8 δ18O
For the global meteoric water line, this is 10‰. Regionally, this parameter varies from much less than 10 to
over 20‰. Indeed, one could ask why it is not 0‰ considering that the origin of precipitation is seawater
with δ18O = δD = 0‰. The reason for this excess is due to kinetic evaporation (non-equilibrium) during the
formation of the primary vapor mass, and it provides an additional tool to trace the origin of water.
Evaporation of water to form atmospheric water vapor is a non-equilibrium or “kinetic” process as it
proceeds at a greater rate in the forward direction than the reverse (condensation). The process becomes
increasingly non-equilibrium with lower humidity. Non-equilibrium evaporation enhances the fractionation
of 18O, and so the vapor is more depleted in this isotope than if the vapor was in equilibrium with the water
(Fig. 4-7). Non-equilibrium evaporation also enhances D fractionation, but not as much. Consequently, the
vapor from which rainout occurs has a slight enrichment in D. The result is the deuterium “excess”
observed in meteoric waters.
When primary evaporation takes place under conditions of extremely low humidity, the deuterium
enrichment is greatest. In the eastern Mediterranean region, the deuterium intercept is 22‰ (Gat and
Carmi, 1970). On average, global freshwaters are formed under conditions of about 85% humidity, and so
global freshwaters plotting on the global meteoric water line have a deuterium excess of only 10‰.
8 Chapter 5 Tracing the Hydrological Cycle
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-60
-20
20
-14 -10 -6 -2 2
δ18O ‰
D ‰
VSMOW
h = 100%
equilibrium
Vapor
Rain
δ2H = 8δ
18O + 0
-100
-60
-20
20
-14 -10 -6 -2 2
δ18O ‰
D ‰
VSMOW
Vapor
Rain
GMWL
δ2H = 8δ
18O + 10
h = 85%
kinetic
evaporation
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-60
-20
20
-14 -10 -6 -2 2
δ18O ‰
D ‰
VSMOW
h = 78%
kinetic
evaporation
δ2H = 8δ
18O + 22
Vapor
Rain
Fig. 4-7 Deuterium excess in primary vapor masses. The lower the ambient humidity during primary evaporation over seawater
increases kinetic evaporation and generates a greater the deuterium excess (d = δD – 8 δ18O), resulting in precipitation that plots on a line with slope near 8 but with a deuterium enrichment. The average deuterium excess for global freshwaters is 10‰, which is the
deuterium intercept for the global meteoric water line.
Local meteoric water lines The slope and the deuterium intercept for the meteoric water line can vary according to humidity conditions during primary evaporation and due to secondary evaporation during rain (in arid regions at least). Any given region will then have its own characteristic meteoric water line, which should be representative of the surface and groundwaters recharged in that region (Fig. 4-5). The IAEA maintains a global network for isotopes in precipitation (GNIP), that can be accessed to gain localized precipitation data and local meteoric water lines (LMWL). These data sets are available through their website < http://isohis.iaea.org/>. In any study of isotope hydrology, it is useful to use precipitation data that best represents the study area. In the absence of regional or local precipitation data, the GMWL is often substituted. However, this can be misleading, particularly in arid and high latitude regions.
CLIMATE EFFECTS ON δ18
O AND δD IN PRECIPITATION
The temperature effect on the isotopic composition of precipitation is one of the primary tools for tracing waters in the hydrologic cycle. As climatic temperature is affected by latitude, altitude, seasonality and
proximity to large water bodies, so too is the δ18O for regional precipitation. These isotope effects are then used to identify areas of recharge for water resources, seasonal aspects of watershed drainage, mixing of different waters. They are even used to reconstruct paleoclimates and paleohydrology.
Continental effect
Continentality is a climate index that incorporates two parameters; the average annual range in temperature
and latitude (Barry and Chorley, 1987). Both influence the isotopic composition of precipitation. The
influence of latitude is observed at the global scale (Fig. 4-4), where δ18Ο gradients increase from broad
equatorial belt of flat gradients to steep gradients nearer the poles. The Mackenzie River, in northwestern
Canada, has δ18O = –18‰ where it discharges into the Beaufort Sea (Hitchon and Krouse, 1972), whereas
local precipitation on this northern coast has δ18O of –21‰. The difference is due to a catchment that
ranges over 15 degrees of latitude.
The continentally of a meteorological station also increases with distance from the coast, as interior stations
experience greater seasonal extremes in temperature than coastal stations. Accordingly, the isotopic
Environmental Geochemistry and Isotopes 9
composition of precipitation becomes more depleted with increasing distance from the coast due to cooling
and rainout along storm tracks.
The combined effect of latitude and distance from the coast that comprise continentality are evident in the
map of weighted mean δ18O in precipitation for North America (Fig. 4-8). The δ
18O isopleths tend to
contour the coastlines, while also showing a strong latitudinal gradient. The highest mean annual δ18O
values occur along the southern coastlines of the Gulf of Mexico and California where much of the
continental precipitation originates. Lowest values are found in the northern interior of the continent where
the rainout process has evolved to low residual vapor fractions, f.
Fig. 4-8 Continental effect of δ18O of precipitation in North America. Based on mean annual values weighted by the amount of
precipitation.
Altitude effect
The correlation between δ18O and temperature provides a useful tool to identify the recharge elevation of
groundwater, given that air temperatures are cooler at higher elevation. Tropospheric temperatures decrease
by about 0.6° per 100-m rise in altitude (the tropospheric lapse rate). From Fig. 4-1, precipitation should
then experience depletion in 18O of about 0.4‰ per 100 m rise in altitude. In fact, the δ18O of precipitation
decreases anywhere from 0.15‰ to 0.5‰ per 100-m rise in altitude, while δD decreases by between –1 to –
4‰. The effect is observed even in watersheds with elevation contrasts of less than a few hundred meters,
provided that sufficient data are collected to resolve seasonal effects and even-scale noise.
Precipitation for catchment in the maritime piedmont of the Italian Alps was sampled at different
elevations, providing the good correlation observed in (Fig. 4-9). In this case, the altitude effect was –
0.31‰ for δ18O per 100 m rise. For Mont Blanc (Moser and Stichler, 1970), the gradient was a much
steeper –4‰ for δD (~–0.5‰ for δ18O) per 100-m rise. Warmer regions tend to have lower temperature
gradients, and so the altitude effect is generally less than –0.2‰ for δ18O per 100–m rise.
Non-alpine catchments with relief contrasts as low as a couple hundred meters also demonstrate an altitude
effect for 18O and D. It is seldom evident from collecting precipitation from one precipitation event.
However, long-term weighted averages of precipitation can resolve a depletion trend with altitude. This is
worthwhile to monitor in a study site given that groundwater recharge itself is close to a long-term
weighted average of precipitation in the recharge area. Accordingly, recharge elevations can even be
distinguished in low-relief catchments.
10 Chapter 5 Tracing the Hydrological Cycle
0
500
1000
1500
2000
2500
-16 -14 -12 -10 -8δ
18O ‰
Ele
vation (
m a
.s.l.)
W Val Corsaglia, Italy
0.32‰ per 100 m
0
100
200
300
400
-0.3 0.2 0.7 1.2δ
18O ‰
Southern Oman Monsoon
0.25‰ per 100 m
Fig. 4-9 The altitude effect for δ18O in precipitation at three sites. Val Corsaglia, a high altitude site in the Italian Alps sampled
during October, 1974 (Bortolami, 1978), the summer monsoon in the Dhofar Mountains of Southern Oman monitored weekly over the 1985 monsoon period (Clark et al.., 1987), and a low relief agricultural catchment in Eastern Ontario monitored over a 15 month period (amount-weighted, Suchy, 2008). The studies for the Oman and the Raisin River sites demonstrate that the depletion trend with
elevation is apparent even for low-relief catchments, providing multiple events or long-term monitoring data are collected.
Seasonal effects in continental precipitation
For temperate and continental regions, a principal control on the isotopic composition of precipitation is the
seasonal change in temperature and consequently in the δ18O and δD of precipitation. The amplitude of this
seasonal effect increases with the continentality (latitude and distance inland) of the site. These seasonal
variations in δ18O and δD give us an important tool to determine watershed response to precipitation, mean
residence time of water in drainage basins, and to monitor groundwater recharge.
Fig. 4-10 shows the seasonal variation for δ18O for stations from Antarctica to northern Greenland. The
greatest seasonal change is observed for sites with high continentality and high latitude (e.g. Edmonton and
Thule). Marine (St. Helena) and coastal sites (Cape Grim) show the least amount of seasonal variation..
The importance of temperature in this correlation is shown in Fig. 4-11. For the continental stations, a
change in temperature of about 1° C changes the average δ18O by about 0.3 to 0.5‰, or about a 2 to 4°C
change in temperature per 1‰ shift in δ18O. For the marine stations at Cape Grim and St. Helena, there is
very little correlation with temperature. This reflects the lack of evolution in the vapor mass before rainout
occurs at these stations.
Unlike the continental stations, the seasonal variation δ18O for many tropical marine stations which receive
monsoon precipitation correlates poorly with temperature. Monsoon rains occur seasonally in tropical
regions when the low altitude trade winds shift inland, bringing moist air inland along continental margins.
As seasonal ∆T is low there is little variation in δ18O and δD. There is also a stronger influence of the
amount effect (Fig. 4-6) in these regions where precipitation during the low-rainfall months experiences
evaporation in the low humidity air column.
Environmental Geochemistry and Isotopes 11
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-15
-10
-5
0
Cape Grim, Tasmania – 41°S
St. Helena, South Atlantic – 16°S
Vernadsky, West Antartica - 65°S Coyhaique, Chile – 45°S
Chicago, Illinois –42°N
Edmonton, Alberta –54°N
Ottaw a, Ontario –45°N
Thule, Greenland –77°N
18O
‰ V
SM
OW
Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec
Fig. 4-10 The seasonal variation in δ18O in precipitation at stations from low to high latitude in the northern and southern
hemispheres. Data are for monthly averages. Note the lack of seasonal effect for the oceanic station – St. Helena, an atoll in the south
Atlantic Ocean, and the subdued seasonal effect for the two coastal stations at Cape Grim, Tasmania, and Vernadsky, a Russian research station near the tip of the West Antarctic peninsula. Strong continentality imparts a high amplitude variation for 18O in precipitation at the inland stations – Chicago, Ottawa and Edmonton.
Edmonton, Alberta18O = 0.47 T - 21
Coyhaique, Chile18O = 0.43 T - 15
Thule, Greenland18O = 0.41 T - 19
Vernadsky, West Antarctica18O = 0.45 T - 9
Chicago, Illinois18O = 0.32 T - 10
Ottaw a, Ontario18O = 0.30 T - 13
-30
-25
-20
-15
-10
-5
0
-35 -25 -15 -5 5 15 25 35
Mean monthly air temperature (°C)
δ1
8O
‰ (
mean m
onth
ly)
Cape Grim, Tasmania
St. Helena, South Atlantic
Fig. 4-11 The correlation between mean monthly air temperature and δ18O in precipitation at stations shown in Fig. 4-10.
12 Chapter 5 Tracing the Hydrological Cycle
Paleoclimate effects
Long-term trends in mean annual temperature can affect the stable isotope contents of precipitation. Direct
monitoring of such trends over the past 30 years shows strong T–δ18O correlation but with amplitudes on
the order of 1 to 2‰ and 1 to 2°C (Rozanski et al., 1993). Long term climate records are available in Arctic
and Antarctic glacier ice cores. The δ18O variations in these cores provide a record of temperature change
that now extends back over some 650,000 years. Core extracted from the Antarctic ice sheet provides
perhaps the most dramatic example of such detailed temperature records.
-450
-425
-400
-375
-350
0 100,000 200,000 300,000 400,000 500,000 600,000 700,000
Years before present (yr BP)
D‰
Warmer
climate
Colder
climate
Holocene
interglacial
Last
glacial
maximum
Fig. 4-12 Change in climate as recorded by the δD in glacier ice from Antarctica (EPICA, 2004). Periods characterized by depleted
values for δD correspond with advances of continental ice sheets. Peak of enriched deuterium in the ice represent the warmer climates
of the interglacial periods.
ISOTOPE EFFECTS OF EVAPORATION
During precipitation and runoff, fresh waters generally retain the δ18O – δD “signature” that reflects their origin. However, if the freshwaters undergo evaporation, then they will become isotopically enriched. This
is because of the loss of light isotopes in the vapor fraction (recall that ε18Ovapor-water is about –10‰).
Further, if the evaporation occurs under conditions of low humidity (i.e. lower than 100% saturation), then non-equilibrium or “kinetic” effects impart an additional fractionation. Such kinetic fractionation is
different for δ18O and δD. The result is that the waters lose their meteoric relationship of slope ~8 on a
δ18O-δD diagram, and plot along a distinct “evaporation slope” typically less than 5. Kinetic effects
increase with lower ambient humidity. If evaporation is minor, then the remaining water may experience no measurable change. If a the water loss
is more than a few percent, the result is a positive shift in the δ18O and δD composition of the waters away from a position on the local meteoric water line (Fig. 4-13). Typically such kinetic evaporation occurs during overland flow in arid landscapes, in lakes and reservoirs, and from soils during slow infiltration of recharge. It can also occur for light rains falling through warm dry air, noted by Dansgaard (1964) as the amount effect.
Environmental Geochemistry and Isotopes 13
-120
-80
-40
0
-16 -12 -8 -4 0 4 8
Precipitation
Surface waters
δ18O ‰ VSMOW
D ‰
VS
MO
W W
Meteoric w ater line
δ2H = 8 δ18O + 10
Evaporation trend
slope = 4.4
average of
precipitation
Fig. 4-13 Evaporation effect observed in surface waters during a dry period. Both 18O and D become enriched in water during evaporation, although this kinetic effect is greater for 18O, resulting in a shift away from the meteoric water line.
Evaporation from surface waters Evaporation is a kinetic fractionation process that favors the light isotopes 16O and H. The heavy isotopes become concentrated in the water phase. If the amount of evaporation is significant compared with the volume of water being evaporated, an isotopic enrichment can be measured in the water. In fact, the
enrichment trend for both δ18O and δD follows a Rayleigh distillation, like rainout (Fig. 4-3). Fig. 4-14
shows this exponential enrichment for 18O and D as the residual water fraction f approaches 0. On a δ18O –
δD diagram, evaporated waters form a positive trend that deviates from the meteoric water line towards 18O enrichment (Fig. 4-14). Evaporative trends have a slope between 2 and 5, depending on the relative humidity.
-50
0
50
100
150
200
00.20.40.60.81
f
-10
0
10
20
30
18O
‰
VS
MO
W W
Original
w ater
D ‰
V
SM
OW
W
δ18O
δD
-50
0
50
100
150
200
-10 0 10 20 30
δD = 8 δ18O + 10
δ18O ‰ VSMOW W
δD = 5 δ18O – 6.2
f =0.2
0.4
0.6
0.8
Fig. 4-14 The isotopic effect of evaporation on water. Left diagram shows the Rayleigh enrichment of 18O and D as the residual fraction of water f approaches 0, for humidity of 50% and at 25°C. The diagram on the right shows the combined evaporation effect
for δ18O and δD as f approaches 0.
The reason for this deviation is found in the kinetic fractionation factors for 18O and D. Gonfiantini (1986) determined empirical relationships based on humidity that is used to determine the isotopic fractionation between water and vapor during non-equilibrium evaporation from surface waters:
δ18Ow – δ18Ov = ε18Ow-v + 14.2 (1 – h)
δDw – δDv = εDw-v + 12.5 (1 – h)
14 Chapter 5 Tracing the Hydrological Cycle
From these two equations, we can see that the additional enrichment in the water is greater for 18O than for D, hence the deviation from the MWL. Further, the influence of humidity h is such that under extremely dry conditions, the slope of the evaporation trend is lower. Gonfiantini (1986) found the following
correlation for the δ18O - δD relationship in evaporating surface waters:
Slope 3.9 4.2 4.5 5.2 6.8 8
Humidity 0 25% 50% 75% 95% 100%
Understanding the isotope effects during evaporation allows us to calculate the amount of water that has been lost to evaporation from a reservoir. This is particularly useful for inventories of water resources in arid regions and for evaluating the efficiency of surface water impoundments.
Example 4-2 Calculating water lost by evaporation from a surface water body.
The evolution of δ18O and δD in water retained by a dam on a dry river course in an arid region has been
monitored to assess losses to evaporation prior to infiltration. As wind mixing of the water column is
efficient, the data are representative of the full water column, which have T = 25°C. What is the
evaporative loss as a fraction of the original water volume and what is the percent infiltration?
A plot of the isotope data shows that these waters have evaporated with a δ18O - δD slope of 4.5. From the above data for slope vs. humidity, This correlates with a humidity of 50%.
s = 4.5
- 60
- 40
- 20
0
20
- 10 - 8 - 6 - 4 - 2 0 2
δ18O
δ2H
GMWL
∆18O = 6.4‰
Original rain
The kinetic enrichment factor for 18O is determined using the empirical relationship introduced above, and
the equilibrium enrichment factor ε18Ov–w = –9.3‰ () (expressing enrichment factor as product–reactant):
δ18Ov – δ18Ow = ∆18Ov–w = ε18Ov–w – 14.2 (1 – h)
= –9.3 – 14.2(1–0.5) = –16.4 ‰ The overall enrichment for 18O under these conditions is –16.4‰. The fraction water loss is calculated from
the difference between the initial and final 18O content of the reservoir water (δ18Ofinal – δ18Oinitial = 6.4‰, from diagram) using the simplified Rayleigh equation.:
δ18Ofinal – δ18Oinitial = ∆18Ototal · lnf = 6.4‰
–16.4 · lnf = 6.4 lnf = –0.39 f = 0.68
This gives a residual water fraction f of 0.68 and so an evaporative loss of 32%. Thus, only 68‰ of the original water volume in the reservoir has infiltrated to the aquifer. Note that this solution is a first-order approximation. A more accurate analysis requires consideration of the isotope value for the ambient water
vapor in the atmosphere, as this back-exchanges with the reservoir water during evaporation.
Environmental Geochemistry and Isotopes 15
Evaporation and transpiration of groundwaters Groundwaters may have an evaporated signal due to evaporation of surface waters prior to infiltration, or due to evaporation of soil moisture from the unsaturated zone. In the latter case, the evaporation effect is even stronger and results in an isotope shift along a line with much lower slope than for surface water evaporation. Allison (1982) used sand columns to show that kinetic evaporation effects increase as soil moisture is lowered by evaporation from the unsaturated zone. The net effect is that between infiltration events, the soil moisture becomes highly enriched and affects the overall isotopic composition of the next recharge event. In the extreme case, evaporation from the water table occurs by capillary flux and vapor diffusion. Under such conditions, the groundwaters have an evaporative enrichment with a much lower slope (Fig. 4-15).
evaporation slope = 2.3
-30
-20
-10
0
10
-5 -4 -3 -2 -1 0 1 2 3 4
MWL
δ18O ‰ VSMOW
D ‰
VS
MO
W
W
Fig. 4-15 Evaporation trend for groundwaters recharged by infiltration through a sand desert. The low slope is due to strong evaporation from the soil column, which enhances the kinetic fractionation effects.
Although groundwater losses by evaporation can be significant in arid regions, the vegetative cover in most other regions precludes significant evaporation from the ground surface. However, forested landscapes and grasslands transpire groundwaters at even greater rates. Over 100 times more H2O is transpired from leaves than the CO2 that is fixed during photosynthesis. A single acacia tree can transpire up to a cubic metre of water per day. However, this transpired water is taken up by the roots in a quantitative manner, with essentially no isotope fractionation. Such losses of soil water by transpiration leaves no isotopic signature on the residual water, although the process does increase soil water salinity.. This is useful in attributing water losses to evaporation and transpiration. Both result in increased salinity, but only evaporation will enrich the soil water in 18O and D. In the case of irrigated agriculture, identifying water losses using isotopes to compliment geochemistry allows more efficient irrigation methods to be developed. Similarly, determining the cause of salt buildup in soils (e.g. Simpson et al., 1987) is the first step in developing remedial plans.
TRACING GROUNDWATER ORIGIN WITH ISOTOPES – A FEW EXAMPLES
The field of isotope hydrology is loaded with case studies for the origin of recharge as well as for groundwater flow and mixing. These are best constrained with the conjunctive use of both isotopes and geochemistry. However, the following are examples where the isotope signature alone of different waters is diagnostic of recharge origin and mixing. Typically, they are based on characteristic trends observed on a
δ18O vs. δD diagram.
16 Chapter 5 Tracing the Hydrological Cycle
Recharge of deep artesian groundwaters in Northern Jordan The karstified Cretaceous limestones of northern Jordan host remarkable artesian groundwaters that supply a significant fraction of the water supply to the capital, Amman. Considering their importance, understanding their recharge source and whether they are part of the highlands of northern Jordan receives precipitation
Fig. 4-16 The Ajoun highlands of northern Jordan with locations of the well fields supplying the capital. (from Google Maps)
-55
-45
-35
-25
-15
-8 -7 -6 -5 -4
δ18O ‰
D ‰ Ajloun fm
groundw aters
Baqla fm
groundw aters
Northern Jordan Meteoric Water Line
δD = 7.0 δ18O + 16
Pleistocene groundwater
Irbid
618 m aslRas Munif
1150 m asl
Deir Alla
190 m asl
Fig. 4-17 Precipitation and artesian groundwaters in the karstic aquifers of northern Jordan. The three precipitation stations are Ras Munif at the summit of the Ajloun highlands in northern Jordan, Irbid on the flank of the uplifted anticline, and Deir Alla located in the Jordan Rift valley to the west of the anticline. Note the significant difference between the actively recharged groundwaters (Ajloun
fm and Baqla fm) and the Pleistocene groundwater from a deeper sandstone aquifer recharged during cooler, wetter conditions over 10,000 years ago.
Ras Munif 1150 m asl Irbid 618 m
Deir Alla 190 m Well fields
Environmental Geochemistry and Isotopes 17
Geothermal waters in Western Canada The recharge origin of hotsprings is a fascinating application of isotopes, particularly because of the strong relief that usually characterizes the catchments where geothermal waters are found. The contrast in elevation that provides the driving mechanism for deep circulation and heating of meteoric waters is manifested by depletion in 18O and D. For both geothermal settings in Fig. 4-18, the hotspring waters are more depleted than the local runoff and so are recharged at the highest elevations in their respective catchments.
-90
-70
-50
-30
-12 -10 -8 -6
δ18
O ‰ VSMOW
δD
‰ V
SM
OW
West Coast Hotsprings
Hotsprings
local runoff
-160
-140
-120
-100
-20 -18 -16 -14
δ18
O ‰ VSMOW
δD
‰ V
SM
OW
Selkirk Range Hotsprings
Hotsprings
local runoff
Fig. 4-18 δ18O vs. δD of hotspring waters compared with local runoff. From the association with the global water line, the thermal
waters clearly originate as meteoric waters in the region. The depletion in the hotspring waters compared to local runoff shows that they are recharged at a high elevation in the catchment. The coastal hotsprings (left chart) are much more enriched than the interior
(Selkirk Range hotsprings, right chart) demonstrating the continental effect where precipitation in the interior is more evolved and hence, more depleted.
REFERENCES AND FURTHER READING
Bortolami, G.C., Ricci, B., Susella, G.F. and Zuppi, G.M., 1979. Isotope hydrology of the Val Corsaglia,
Maritime Alps, Piedmont, Italy. In: Isotope Hydrology 1978, Vol. I, IAEA Symposium 228, June 1978, Neuherberg, Germany: 327-350.
Craig, H., 1961. Isotopic variations in meteoric waters. Science, 133: 1702-1703. Craig published his
remarkable findings on the correlation between 18O and D in global fresh waters in 1961. This is the basis
of the meteoric relationship which defines the global and local meteoric water lines.
Dansgaard, W., 1964. Stable isotopes in precipitation. Tellus, 16: 436-468. This is the original publication
on the 18O – temperature effect in precipitation, and is the basis of isotope hydrology today.
Clark, I.D. and Fritz, P., 1997. Environmental Isotopes in Hydrogeology, Lewis Publishers, Boca Raton, FL. Chapters 2 and 3, pp. 36-78. EPICA Community Members, 2004. Eight glacial cycles from an Antarctic ice core, Nature, 429 (6992), 623-628. IAEA, 1981. Stable Isotope Hydrology – Deuterium and Oxygen-18 in the Water Cycle. J.R. Gat and R. Gonfiantini (eds.), Technical Report Series No. 210, International Atomic Energy Agency, Vienna.
18 Chapter 5 Tracing the Hydrological Cycle
Moser, H. and Stichler, W., 1970. Deuterium measurements on snow samples from the Alps. Isotope
Hydrology 1970, IAEA Symposium 129, March, 1970, Vienna: 43-57.