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Old groundwaters

Old groundwaters. Istv án 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

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Old groundwaters

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  1. Old groundwaters István Fórizs Ph.D. Institute for Geochemical Research, Hungarian Academy of Sciences Budapest

  2. 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

  3. Nomenclature • Old groundwaters are • Paleo-groundwaters (older than 10 000 a, infiltrated during the latest glaciation) • Sub-modern (older than 60 a)

  4. Stable isotopes and paleo-groundwaters • These waters were infiltrated at cooler climatic conditions during the Ice Age. • Their dD and d18O 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.

  5. Example: Oman

  6. Shift in deuterim-excess (d-excess) • Effect of primary evaporation • Effect of secondary evaporation • Definition: d = dD – 8*d18O

  7. Effect of relative humidity (h) of the air:Primary evaporation Sea water 50% 85% 100% Global Meteoric Water Line

  8. Secondary evaporation 20% 100% 40% 60% 80% 40% 80% 60% 20% Initial water (lake or rain drop) GMWL

  9. Continental effect vapour vapour vapour rain rain Sea Continent d18O

  10. (Triassic) Bunter sandstone, England Bath et al. 1979

  11. Ice cores show well the climate change

  12. GISP2Ice core,Greenland Age (year)

  13. Ice cores: Canada, Greenland, Antarctic

  14. Chemistry and paleo-groundwaters

  15. Conceptual model of groundwater flow

  16. 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

  17. Groundwater dating methods

  18. Groundwater dating methods • Radiocarbon: 14C • Chlorine-36: 36Cl • The uranium decay series • Helium ingrowth • Krypton-81: 81Kr

  19. 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-lt • A0 and At are 14C initial activity, and activity after time ‘t’, l is decay constant.

  20. Rearranged decay equation t = -8267×ln(At/A0) [year]

  21. T1/2: Half-life Aoinitial activity

  22. 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%.

  23. Source of 14C • Natural:147N + 10n → 146C + 11p • Where n = neutron, p = proton • Anthropogenic: nuclear bomb tests starting in 1952.

  24. Natural variation in atmospheric 14C

  25. 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.

  26. Anthropogenic impacts on atmospheric 14C

  27. 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.

  28. Decay equation becomes: At = qA0e-lt or t = -8267×ln(At/(qA0)) [year]

  29. Short introduction to carbon stable isotope geochemistry

  30. Abundance of carbon stable isotopes 12C = 98,9% 13C = 1,1%

  31. 13C distribution in nature

  32. 13C in C3, C4and CAM plants

  33. Photosinthesis • C3plants (85%): Calvin cycle E.g.trees, cereals, legumes (bean), beet. • C3plants:d13C value is from -33to -20 [‰]VPDB • Mean value= -27‰.

  34. Photosinthesis • C4plants (5%): Hatch-Slack cycle E.g.cane, maize • d13C value is -16to -9 [‰]VPDB • Mean value:-12,5‰.

  35. 13C in soil CO2 • Soil CO2originates 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 CO2diffuses to the atmosphere causing isotopic fractionation: the remaining CO2 is heavier by ca. 4‰. • The d13C value of soil CO2: C3vegetation: ≈ -23 [‰]VPDB C4vegetation:≈ -9 [‰]VPDB

  36. Carbon in water • Source: air CO2 (d13C ≈ -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) • CO32- (carbonate ion) } DIC

  37. Distribution of carbonate species as a function of pH at 25 °CClark-Fritz 1997

  38. Isotopic fractionation at 25 °C • Soil CO2 • CO2(aq) • H2CO3 • HCO3- • CO32- } εCO2(aq)-CO2(g) = -1.1‰ } CO2(aq)≡ H2CO3 } εHCO3(-)-CO2(aq) = 9.0‰ } εCO3(2-)-HCO3(-) = -0.4‰

  39. 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

  40. Fractionation: 25 °C, DIC-CO2(soil)Clark-Fritz 1997

  41. Fractionation: DIC-CO2(soil) at 25 °CClark-Fritz 1997

  42. The pathway of 14C to groundwater in the recharge environment

  43. Correction methods • Statistical • Chemical mass-balance • d13C • Dolomite dissolution • Matrix exchange (Fontes-Garnier model)

  44. 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

  45. 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.

  46. Chemical mass-balance 2 • Calculation by chemical data mDICfinal = mDICrecharge+[mCa2+ + mMg2+ -mSO42- + ½(mNa+ + mK+ - mCl-)] m = concentration in moles/liter

  47. d13C mixing model 1 • Closed system model at low pH d13Csample - d13Ccarb q = ───────────────, d13Csoil CO2 - d13Ccarb Where d13Csample = measured in groundwater DIC d13Ccarb = 0 ‰ (calcite being dissolved) d13Csoil CO2 = -23‰

  48. d13C mixing model 2 • Closed system model at any pH d13Csample - d13Ccarb q = ───────────────, d13Crecharge - d13Ccarb Where d13Crecharge = d13Csoil CO2+ e13CDIC-CO2(soil)

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