Old groundwaters

István Fórizs Ph.D.Institute for Geochemical Research,

Hungarian Academy of SciencesBudapest

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 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.

Example: Oman

Shift in deuterim-excess (d-excess)

• Effect of primary evaporation

• Effect of secondary evaporation

• Definition: d = D – 8*18O

-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

-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)

Continental effectContinental effect

18OSea

Continent

vapour

vapour vapour

rainrain

(Triassic) Bunter sandstone, EnglandBath et al. 1979

-120

-110

-100

-90

-80

-70

-60

-16 -15 -14 -13 -12 -11 -10

δ18O

δ2H

GMWL SPRING RIVER BORE

Ice cores show well the climate change

GISP2Ice core,

Greenland

0

5000

10000

15000

20000

25000

30000

-45 -40 -35 -30 -25

18O [‰]VSMOW

kor

[év]

Age

(ye

ar)

Ice cores: Canada, Greenland, Antarctic

Chemistry and paleo-groundwaters

Conceptual model of groundwater flow

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

Groundwater dating methods

Groundwater dating methods

• Radiocarbon: 14C

• Chlorine-36: 36Cl

• The uranium decay series

• Helium ingrowth

• Krypton-81: 81Kr

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.

Rearranged decay equation

t = -8267×ln(At/A0) [year]

T1/2: Half-life

Ao initial activity

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%.

Source of 14C

• Natural: 147N + 1

0n → 146C + 1

1p

• Where n = neutron, p = proton

• Anthropogenic: nuclear bomb tests starting in 1952.

Natural variation in atmospheric 14C

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.

Anthropogenic impacts on atmospheric 14C

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.

• Decay equation becomes:

At = qA0e-t

or

t = -8267×ln(At/(qA0)) [year]

Short introduction to carbon stable isotope geochemistry

Abundance of carbon stable isotopes

12C = 98,9%13C = 1,1%

13C distribution in nature

13C in C3, C4 and CAM plants

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‰.

Photosinthesis

• C4 plants (5%): Hatch-Slack cycle

E.g. cane, maize

• 13C value is -16 to -9 [‰]VPDB

• Mean value: -12,5‰.

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

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

Distribution of carbonate species as a function of pH at 25 °C

Clark-Fritz 1997

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‰

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

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

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

The pathway of 14C to groundwater in the recharge environment

Correction methods

• Statistical

• Chemical mass-balance• 13C

• 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 – 0.75 for karst systems

• 0.75 – 0.90 for sediments with fine-grained carbonate such as loess

• 0.90 – 1.00 for crystalline rocks

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.

Chemical mass-balance 2

• Calculation by chemical data

mDICfinal = mDICrecharge +[mCa2+ + mMg2+ -mSO4

2- + ½(mNa+ + mK+ - mCl-)]

m = concentration in moles/liter

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 ‰

13C mixing model 2

• Closed system model at any pH

13Csample - 13Ccarb

q = ───────────────,

13Crecharge - 13Ccarb

Where

13Crecharge = 13Csoil CO2 + 13CDIC-CO2(soil)

: enrichment factor

• Depends highly on pH and on temperature

13CA-B = (RA / RB - 1)*1000 ‰,

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)

• mDICCO2-exch = (13CmeasxmDICmeas - 13CcarbxmDICcarb -

13Csoilx(mDICmeas – mDICcarb)/(13Csoil - 13CCO2(soil)-CaCO3 -

13Ccarb)

• this may be negative

• qF-G = (mDICmeas – mDICcarb + mDICCO2-exch)/ mDICmeas

Uncertainity

(Triassic) Bunter sandstone, EnglandBath et al. 1979

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?

Clorine-36: 36Cl

Chlorine isotopes

35Cl = 75.4% stable36Cl = radioactive, 301 000 year half-life

37Cl = 24.6% stable

Sources of 36Cl

• Natural: collision of cosmic neutron and 35Cl atom.

• Subsurface or epigenic production?

• Anthropogene: mostly nuclear bomb tests in sea water.

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

Decay

At = A0e-t

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

• 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

Krypton-81: 81Kr

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.

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

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

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)

Brines