old groundwaters istván fórizs ph.d. institute for geochemical research, hungarian academy of...
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Old groundwaters
István Fórizs Ph.D.Institute for Geochemical Research,
Hungarian Academy of SciencesBudapest
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Why should we identify old groundwaters?
• To determine the time and place of recharge (recharge may already be stopped)
• Mean residence time
• Exploitation induced recharge
• To understand the geochemical and hydrological processes
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Nomenclature
• Old groundwaters are• Paleo-groundwaters (older than 10 000 a,
infiltrated during the latest glaciation)• Sub-modern (older than 60 a)
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Stable isotopes and paleo-groundwaters
• These waters were infiltrated at cooler climatic conditions during the Ice Age.
• Their D and 18O values are significantly more negative than those of Holocene infiltrated ones. Temperature effect!!
• Shift in d-excess. The effect of relative humidity of (h) air on the primary evaporation. Characteristic for arid regions, Eastern Mediterranean and North Africa.
• There are some areas where paleo-groundwaters post-date the glaciation, because during the Ice Age there was a permanent ice cover. The melted water infiltrated during the deglaciation (early Holocene), e.g. in Canada.
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Example: Oman
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Shift in deuterim-excess (d-excess)
• Effect of primary evaporation
• Effect of secondary evaporation
• Definition: d = D – 8*18O
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-140
-120
-100
-80
-60
-40
-20
0
20
40
-18 -16 -14 -12 -10 -8 -6 -4 -2 0 2 4
18O [‰]
D [
‰]
Effect of relative humidity (h) of the air:Primary evaporation
Global Meteoric Water Line
100%85%
50%
Sea water
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-140
-120
-100
-80
-60
-40
-20
0
20
40
-18 -16 -14 -12 -10 -8 -6 -4 -2 0 2 4
18O [‰]
D [
‰]
Secondary evaporation
GMWL
20%
40%60%
80%
20%
40%60%
80%
100%
Initial water (lake or rain drop)
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Continental effectContinental effect
18OSea
Continent
vapour
vapour vapour
rainrain
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(Triassic) Bunter sandstone, EnglandBath et al. 1979
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-120
-110
-100
-90
-80
-70
-60
-16 -15 -14 -13 -12 -11 -10
δ18O
δ2H
GMWL SPRING RIVER BORE
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Ice cores show well the climate change
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GISP2Ice core,
Greenland
0
5000
10000
15000
20000
25000
30000
-45 -40 -35 -30 -25
18O [‰]VSMOW
kor
[év]
Age
(ye
ar)
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Ice cores: Canada, Greenland, Antarctic
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Chemistry and paleo-groundwaters
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Conceptual model of groundwater flow
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Chemistry and paleo-groundwaters
• Water-rock interaction may change the chemistry of water significatly
• Recharge area:– low TDS– frequently Ca-HCO3 type
• Discharge area:– high TDS– frequently Na(-Ca)-HCO3(-Cl-SO4) type– high pH– high trace element content
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Groundwater dating methods
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Groundwater dating methods
• Radiocarbon: 14C
• Chlorine-36: 36Cl
• The uranium decay series
• Helium ingrowth
• Krypton-81: 81Kr
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Basis of 14C age determination
• Radioactive decay (discovered by Libby in 1946, Nobel Prize).
• Half-life of 14C is 5730 a (years).
• Decay equation:
At = A0×e-t
• A0 and At are 14C initial activity, and activity after time ‘t’, is decay constant.
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Rearranged decay equation
t = -8267×ln(At/A0) [year]
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T1/2: Half-life
Ao initial activity
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Expression of 14C activity
• 14C is expressed versus a reference, in percent modern carbon, pmC.
• Reference is the pre-industrial 14C activity of atmospheric CO2, that is regarded as 100%.
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Source of 14C
• Natural: 147N + 1
0n → 146C + 1
1p
• Where n = neutron, p = proton
• Anthropogenic: nuclear bomb tests starting in 1952.
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Natural variation in atmospheric 14C
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The calculated age
• If we disregard the natural variation in atmospheric 14C (A0 is regarded to have been constant, as 100%), then the calculated age is radiocarbon years and not in calendar years.
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Anthropogenic impacts on atmospheric 14C
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Correction: why needed?
• During the flow path 14C is diluted by geochemical reactions:
– Limestone (calcite) dissolution
– Dolomite dissolution
– Exchange with the aquifer matrix
– Oxidation of old organics within the aquifer
• Calcite, dolomite and old organics are free of 14C.
• Initial 14C activity: Arecharge = q* A0,
where q is dilution factor.
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• Decay equation becomes:
At = qA0e-t
or
t = -8267×ln(At/(qA0)) [year]
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Short introduction to carbon stable isotope geochemistry
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Abundance of carbon stable isotopes
12C = 98,9%13C = 1,1%
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13C distribution in nature
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13C in C3, C4 and CAM plants
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Photosinthesis
• C3 plants (85%): Calvin cycle
E.g. trees, cereals, legumes (bean), beet.
• C3 plants: 13C value is from -33 to -20 [‰]VPDB
• Mean value= -27‰.
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Photosinthesis
• C4 plants (5%): Hatch-Slack cycle
E.g. cane, maize
• 13C value is -16 to -9 [‰]VPDB
• Mean value: -12,5‰.
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13C in soil CO2
• Soil CO2 originates from decomposition of organic material and root respiration.
• The pressure of soil CO2 gas is 10-100 times higher than the atmospheric .
• A part of soil CO2 diffuses to the atmosphere causing isotopic fractionation: the remaining CO2 is heavier by ca. 4‰.
• The 13C value of soil CO2:
C3 vegetation: ≈ -23 [‰]VPDB
C4 vegetation: ≈ -9 [‰]VPDB
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Carbon in water
• Source: air CO2 (13C ≈ -7 [‰]VPDB), or soil CO2 ( -9‰ — -23‰) or limestone (0±2‰)
Carbonate species in water• CO2(aq) (aquatic carbondioxide)• H2CO3 (carbonic acid)• HCO3
- (bicarbonate ion)• CO3
2- (carbonate ion)
}DIC
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Distribution of carbonate species as a function of pH at 25 °C
Clark-Fritz 1997
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Isotopic fractionation at 25 °C
• Soil CO2
• CO2(aq)
• H2CO3
• HCO3-
• CO32-
} CO2(aq) ≡ H2CO3
}
}
}
εCO2(aq)-CO2(g) = -1.1‰
εHCO3(-)-CO2(aq) = 9.0‰
εCO3(2-)-HCO3(-) = -0.4‰
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Fractionation factors as a function of temperature
• 103 lnα13CCO2(aq)-CO2(g) = -0.373(103T-1) + 0.19
• 103 lnα13CHCO3(-)-CO2(g) = 9.552(103T-1) + 24.10
• 103 lnα13CCO3(2-)-CO2(g)= 0.87(103T-1) + 3.4
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Fractionation: 25 °C, DIC-CO2(soil)Clark-Fritz 1997
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Fractionation: DIC-CO2(soil) at 25 °CClark-Fritz 1997
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The pathway of 14C to groundwater in the recharge environment
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Correction methods
• Statistical
• Chemical mass-balance• 13C
• Dolomite dissolution
• Matrix exchange (Fontes-Garnier model)
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Statistical model
• If we do not know anything about the recharge area, we can use the world average for q, which is 85% (0.85).
• 0.65 – 0.75 for karst systems
• 0.75 – 0.90 for sediments with fine-grained carbonate such as loess
• 0.90 – 1.00 for crystalline rocks
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Chemical mass-balance• Closed system model: no exchange between DIC
and soil CO2
mDICrecharge
q = ─────────── mDICsample(final)
• m = concentration in moles/liter• mDICrecharge is measured at the recharge area or
calculated from estimated PCO2-pH conditions. If the present climate differs significantly from that during the infiltration, then the calculation is rather speculative.
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Chemical mass-balance 2
• Calculation by chemical data
mDICfinal = mDICrecharge +[mCa2+ + mMg2+ -mSO4
2- + ½(mNa+ + mK+ - mCl-)]
m = concentration in moles/liter
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13C mixing model 1
• Closed system model at low pH
13Csample - 13Ccarb
q = ───────────────,13Csoil CO2 - 13Ccarb
Where13Csample = measured in groundwater DIC
13Ccarb = 0 ‰ (calcite being dissolved)13Csoil CO2 = -23 ‰
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13C mixing model 2
• Closed system model at any pH
13Csample - 13Ccarb
q = ───────────────,
13Crecharge - 13Ccarb
Where
13Crecharge = 13Csoil CO2 + 13CDIC-CO2(soil)
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: enrichment factor
• Depends highly on pH and on temperature
13CA-B = (RA / RB - 1)*1000 ‰,
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Fontes-Garnier model
• Open and closed system dissolution are considered
• mDICcarb = mCa + mMG –mSO4 + ½(mNa + mK –mCl)
• This DIC consists of two parts:• dissolved in open system: C-14 exchange with soil
CO2• dissolved in closed system (C-14 dead)
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• mDICCO2-exch = (13CmeasxmDICmeas - 13CcarbxmDICcarb -
13Csoilx(mDICmeas – mDICcarb)/(13Csoil - 13CCO2(soil)-CaCO3 -
13Ccarb)
• this may be negative
• qF-G = (mDICmeas – mDICcarb + mDICCO2-exch)/ mDICmeas
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Uncertainity
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(Triassic) Bunter sandstone, EnglandBath et al. 1979
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Problem
Data got on well water in Hungary
• Tritium: 3 TU
• 18O = -10,7 [‰]VSMOW
• 14C-content: 30 pmC
• What is your opinion about this water?
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Clorine-36: 36Cl
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Chlorine isotopes
35Cl = 75.4% stable36Cl = radioactive, 301 000 year half-life
37Cl = 24.6% stable
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Sources of 36Cl
• Natural: collision of cosmic neutron and 35Cl atom.
• Subsurface or epigenic production?
• Anthropogene: mostly nuclear bomb tests in sea water.
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Terminology
• R36Cl= number of 36Cl atoms per/Cl
• A36Cl=number of 36Cl atoms/liter
• Evaporation:– R36Cl = constant– A36Cl increase
• Dissolution of „old” chlorine:– R36Cl decrease– A36Cl = constant
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Decay
At = A0e-t
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Initial activity of 36Cl
• A0 is determined by the geomagnetic latitude
• Minimum at 0 and 90 degrees
• Maximum at 40 degrees
• You must take into account the distance from the sea
• You have to create 36Cl/Cl in precipitation map
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• AMS is used for the measurement
• Sampling is very simple
• Geochemical modelling is necessary: dissolution of 36Cl-free chlorine (this is a most problematic part)
• Age range up to 1.5 million years
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Krypton-81: 81Kr
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Krypton-81: 81Kr
• 81Kr is produced in the upper atmosphere by cosmic-ray-induced spallation of five heavier Kr isotopes, i.e. from 82Kr to 86Kr. Or by neutron capture:
8036Kr + n → 81
36Kr + • No significant subsurface production.• No appreciable anthropogenic source.• Half-life is 229 000 years.• Age range: from 35 000 to 670 000 years.
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Krypton-81: 81Kr (cont.)
• The decay equation is:81Krt = 81Kr0×e-t
• The 81Kr concentration is expressed as number of atoms/liter
• 81Kr0 = 1100 atoms/L: initial value in modern groundwater
• E.g. 81Kr = 900 atoms/L
• t = -(ln(900/1100)/ = 66 297 a
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Krypton-81: 81Kr (cont.)
• The 81Kr concentration can be expressed as percent of modern atmosphere (similar to 14C)
• R/Rair = (81Kr/Kr)sample/(81Kr/Kr)air in percent
• E.g. 81Kr = 40%• t = -(ln(40%/100%)/ =
-(ln(0.4)/(3.03*10-6) = 302 722 a
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Krypton-81: 81Kr (cont.)
• Advantages: – Anthropogenic sources are minimal.– 81Kr is inert (no chemical reactions envolved)
• Disadvantages:– Technical difficulties, 1 or 2 labs in the world.– Limited experience (only 3 case studies worldwide)
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Brines
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