Presentation on theme: "Old groundwaters István Fórizs Ph.D. Institute for Geochemical Research, Hungarian Academy of Sciences Budapest."— Presentation transcript:
Old groundwaters István Fórizs Ph.D. Institute for Geochemical Research, Hungarian Academy of Sciences Budapest
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
Nomenclature Old groundwaters are Paleo-groundwaters (older than 10 000 a, infiltrated during the latest glaciation) Sub-modern (older than 60 a)
Stable isotopes and paleo- groundwaters These waters were infiltrated at cooler climatic conditions during the Ice Age. Their D and 18 O 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.
Chemistry and paleo-groundwaters Water-rock interaction may change the chemistry of water significatly Recharge area: –low TDS –frequently Ca-HCO 3 type Discharge area: –high TDS –frequently Na(-Ca)-HCO 3 (-Cl-SO 4 ) type –high pH –high trace element content
Radiocarbon: 14 C Chlorine-36: 36 Cl The uranium decay series Helium ingrowth Krypton-81: 81 Kr
Basis of 14 C age determination Radioactive decay (discovered by Libby in 1946, Nobel Prize). Half-life of 14 C is 5730 a (years). Decay equation: A t = A 0 ×e - t A 0 and A t are 14 C initial activity, and activity after time t, is decay constant.
Rearranged decay equation t = -8267×ln(A t /A 0 ) [year]
The calculated age If we disregard the natural variation in atmospheric 14 C (A 0 is regarded to have been constant, as 100%), then the calculated age is radiocarbon years and not in calendar years.
Correction: why needed? During the flow path 14 C 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 14 C. Initial 14 C activity: A recharge = q* A 0, where q is dilution factor.
Decay equation becomes: A t = qA 0 e - t or t = -8267×ln(A t /(qA 0 )) [year]
Short introduction to carbon stable isotope geochemistry
Abundance of carbon stable isotopes 12 C = 98,9% 13 C = 1,1%
Photosinthesis C 3 plants (85%): Calvin cycle E.g. trees, cereals, legumes (bean), beet. C 3 plants: 13 C value is from -33 to -20  VPDB Mean value= -27.
Photosinthesis C 4 plants (5%): Hatch-Slack cycle E.g. cane, maize 13 C value is -16 to -9  VPDB Mean value: -12,5.
13 C in soil CO 2 Soil CO 2 originates from decomposition of organic material and root respiration. The pressure of soil CO 2 gas is 10-100 times higher than the atmospheric. A part of soil CO 2 diffuses to the atmosphere causing isotopic fractionation: the remaining CO 2 is heavier by ca. 4. The 13 C value of soil CO 2 : C 3 vegetation: -23  VPDB C 4 vegetation: -9  VPDB
Carbon in water Source: air CO 2 ( 13 C -7  VPDB ), or soil CO 2 ( -9 -23) or limestone (0±2) Carbonate species in water CO 2(aq) (aquatic carbondioxide) H 2 CO 3 (carbonic acid) HCO 3 - (bicarbonate ion) CO 3 2- (carbonate ion) } DIC
Distribution of carbonate species as a function of pH at 25 °C Clark-Fritz 1997
Isotopic fractionation at 25 °C Soil CO 2 CO 2(aq) H 2 CO 3 HCO 3 - CO 3 2- } CO 2(aq) H 2 CO 3 } } } ε CO2(aq)-CO2(g) = -1.1 ε HCO3(-)-CO2(aq) = 9.0 ε CO3(2-)-HCO3(-) = -0.4
Fractionation factors as a function of temperature 10 3 lnα 13 C CO2(aq)-CO2(g) = -0.373(10 3 T -1 ) + 0.19 10 3 lnα 13 C HCO3(-)-CO2(g) = 9.552(10 3 T -1 ) + 24.10 10 3 lnα 13 C CO3(2-)-CO2(g) = 0.87(10 3 T -1 ) + 3.4
Fractionation: 25 °C, DIC-CO 2(soil) Clark-Fritz 1997
Fractionation: DIC-CO 2(soil) at 25 °C Clark-Fritz 1997
The pathway of 14 C to groundwater in the recharge environment
Correction methods Statistical Chemical mass-balance 13 C Dolomite dissolution Matrix exchange (Fontes-Garnier model)
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 – 0.75for karst systems 0.75 – 0.90for sediments with fine- grained carbonate such as loess 0.90 – 1.00for crystalline rocks
Chemical mass-balance Closed system model: no exchange between DIC and soil CO 2 mDIC recharge q = mDIC sample(final) m = concentration in moles/liter mDIC recharge is measured at the recharge area or calculated from estimated P CO2 -pH conditions. If the present climate differs significantly from that during the infiltration, then the calculation is rather speculative.
Chemical mass-balance 2 Calculation by chemical data mDIC final = mDIC recharge +[mCa 2+ + mMg 2+ - mSO 4 2- + ½(mNa + + mK + - mCl - )] m = concentration in moles/liter
13 C mixing model 1 Closed system model at low pH 13 C sample - 13 C carb q =, 13 C soil CO2 - 13 C carb Where 13 C sample = measured in groundwater DIC 13 C carb = 0 (calcite being dissolved) 13 C soil CO2 = -23
13 C mixing model 2 Closed system model at any pH 13 C sample - 13 C carb q =, 13 C recharge - 13 C carb Where 13 C recharge = 13 C soil CO2 + 13 C DIC-CO2(soil)
: enrichment factor Depends highly on pH and on temperature 13 C A-B = (R A / R B - 1)*1000,
Fontes-Garnier model Open and closed system dissolution are considered mDIC carb = 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)
mDIC CO2-exch = ( 13 C meas xmDIC meas - 13 C carb xmDIC carb - 13 C soil x(mDIC meas – mDIC carb )/( 13 C soil - 13 C CO2(soil)- CaCO3 - 13 C carb ) this may be negative q F-G = (mDIC meas – mDIC carb + mDIC CO2-exch )/ mDIC meas
Initial activity of 36 Cl A 0 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 36 Cl/Cl in precipitation map
AMS is used for the measurement Sampling is very simple Geochemical modelling is necessary: dissolution of 36 Cl-free chlorine (this is a most problematic part) Age range up to 1.5 million years
81 Kr is produced in the upper atmosphere by cosmic-ray-induced spallation of five heavier Kr isotopes, i.e. from 82 Kr to 86 Kr. Or by neutron capture: 80 36 Kr + n 81 36 Kr + No significant subsurface production. No appreciable anthropogenic source. Half-life is 229 000 years. Age range: from 35 000 to 670 000 years.
Krypton-81: 81 Kr (cont.) The decay equation is: 81 Kr t = 81 Kr 0 ×e - t The 81 Kr concentration is expressed as number of atoms/liter 81 Kr 0 = 1100 atoms/L: initial value in modern groundwater E.g. 81 Kr = 900 atoms/L t = -(ln(900/1100)/ = 66 297 a
Krypton-81: 81 Kr (cont.) The 81 Kr concentration can be expressed as percent of modern atmosphere (similar to 14 C) R/R air = ( 81 Kr/Kr) sample /( 81 Kr/Kr) air in percent E.g. 81 Kr = 40% t = -(ln(40%/100%)/ = - (ln(0.4)/(3.03*10 -6 ) = 302 722 a
Krypton-81: 81 Kr (cont.) Advantages: –Anthropogenic sources are minimal. – 81 Kr 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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