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```DIC para o campo do
bicarbonatobicarbonato
Efetivo para a zona não saturada (sistema Efetivo para a zona não saturada (sistema
aberto) onde mais CO2 pode ser dissolvido aberto) onde mais CO2 pode ser dissolvido
nos solos, mas pouco efetivo para zona nos solos, mas pouco efetivo para zona
No caso da No caso da anortitaanortita (slide25) o Ca(slide25) o Ca2+2+ tbtb atuaatua
Quanto maior o CO2(g) menor será o pH inicial da água
subterrânea, que irá então ser tamponado pelo intemperismo de
solos e rochas. A dissolução da calcita é acelerada pelo ácido
carbônico do solo:
CO2(g) + H20 + CaCO3 → Ca2+ + 2HCO3- (KT = 10-6,41 > Kcalcita)
Essas reações produzem troca iônica e fracionamento isotópico:
Fracionamento do Fracionamento do δδ1313CCCO2CO2 no solono solo
Fatores de fracionamento para espFatores de fracionamento para espéécies de carbonato variam com a cies de carbonato variam com a
temperatura de acordo com as reatemperatura de acordo com as reaçções e as equaões e as equaçções correspondentes ões correspondentes
estão expressas na tabela 1estão expressas na tabela 1
-
14 14 14
2 13
14
13
14
13
14
2
13
14
14
13
2
Dissolução de carbonatos (aqüíferos calcários ou c/ carbonatos secundário) na
área de recarga. Assim, a atividade do 14C no DIC da recarga após a dissolução
(a 14Crec) é igual ao 14C moderno dos solos (a0 14C) vezes o fator de diluição (q):
a 14Crec = q . a0 14C ∴ at 14C = a 14Crec . e-λt ∴ t = -8267 . ln (at 14C / a 14Crec )
EvoluEvoluççãoão geoqugeoquíímicamica do do 1414C no C no
DICDIC
1414COCO22 do solo = do solo = 100 100 pMCpMC
DIDI1414CC
MedidoMedido
DiluiDiluiççãoão do do 1414C C ::
CaCOCaCO33 + + 1414COCO2 2 + H+ H22O O → CaCa2+2+ + H+ H1414COCO33 + HCO+ HCO33 ~~ 50 50 pMCpMC
1414C decayC decay << 50 50 pMCpMC
<<<< 50 50 pMCpMC
DiluiDiluiççãoão vcvc decaimentodecaimento??
MODELOS DE CORREMODELOS DE CORREÇÇÃO PARA ÃO PARA
A DISSOLUA DISSOLUÇÇÃO DE ÃO DE
CARBONATOSCARBONATOS
Statistical correction (STAT model)
Statistical models assume that after the initial carbon uptake in the soil zone by infiltrating water, some
14C dilution will occur through the addition of 14C-free carbon. Statistical evaluations are possible if
geochemical evolution can be averaged over the recharge area to estimate an “initial” value for the 14C
activity of the aqueous carbonate. This initial value represents the fraction of 14C remaining after

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

The lower q values emphasize the importance of carbonate reactions in aquifers with abundant
carbonate whereas in crystalline rocks with little carbonate, q values close to 1 have been determined.
The corrected age is then determined from the decay equation, using this estimate of q:

t 8267 ln C
q C
14
DIC
STAT o
14= − ⋅ ⋅
a
a

The simple statistical approach can be problematic. Geochemical reactions that take place in
groundwaters beyond the recharge area (above) must also be corrected for. The actual 14C dilution is
usually much larger (and thus the q-factors smaller) in older systems than in young ones. Geochemical
models must then be considered.
Alkalinity correction (ALK model)

Where the recharge area cannot be studied, geochemical models can be used that are based on the
parameters measured in the water being dated. The Tamers (1975) or "chemical" correction based on
initial and final carbonate (DIC) concentrations. It was proposed for groundwater in which calcite is
dissolved under closed system conditions:
q
H CO HCO
H CO HCOALK =
+
+
−
−
m m
m m
2 3 3
2 3 3
½
For the vast majority of groundwaters, this leads automatically to a q-factor of about 0.5 because of the
reaction:

H2CO2 + CaCO3 ↔ 2HCO3– + Ca2+

whereby most of the carbonic acid is consumed by limestone dissolution, and the original 14C from soil
carbon dioxide is diluted to about 50%. This model assumes fully closed system conditions, where no
exchange with soil CO2 during calcite dissolution occurs. This model is of limited interest due to its
simplification of geochemical reaction.
CMB-Alk
CMB-Chem
(Mistura com δ 13C)
Matrix exchange (Fontes-Garnier model)

The exchange of carbon isotopes between the DIC and carbonate minerals in the aquifer matrix is often
considered as a cause for δ13CDIC enrichment and 14C dilution. Groundwaters that are essentially at
equilibrium with calcite will exchange carbonate across the mineral-solution interface where CO32– and
Ca2+ are in a continual process of recrystallization. Constraints to such a process include the grain size
of carbonate minerals in the aquifer, the volume and geometry of pores and the degree of “aging” of the
mineral surfaces. The “reactivity” of the mineral surfaces decreases with increased crystallization and
age.

Isotope exchange should be recorded by closed system enrichments to 13C, and so radiocarbon ages can
be corrected according to the 13C mass-balance equation discussed above. However, if secondary
calcite surfaces are involved, then the δ13C will be depleted from that of the host limestone (or
dolomite), and the effect may not be fully accounted for. A more complicated correction is proposed by
Maloszewski and Zuber (1991) who model the retardation of 14C as a sorption reaction. A limitation to
their approach is the sensitivity of the model to distribution coefficient for 14C and pore geometry,
which vary between aquifers.

Fontes and Garnier (1979) developed a correction model (the F-G model) that uses cation
concentrations to determine the contribution of 14C-free matrix carbonate, and isotope mass balance to
apportion 14CDIC into that exchanged with (i) CO2 gas in the soil (open system exchange), and (ii) the
carbonate matrix (matrix exchange). In their model, the total of matrix-derived carbonate is calculated
as:

mDICcarb = mCa2+ + mMg2+ – mSO42– + ½(mNa+ + mK+ – mCl–)

This accounts for carbonate dissolution based on Ca and Mg, with a correction for evaporite dissolution
(mSO42–) and cation exchange (mNa+ + mK+, with a correction for Na+ from salt, mCl–). Their mNO3–
term is dropped here as it is negligible in old waters. This DIC is then apportioned into two components
— that which has exchanged with the soil CO2 (14C active) in an open system and that which has
exchanged with the carbonate matrix (considered to be 14C-dead) under closed system conditions. The
fraction of this DIC that has exchanged with soil CO2 in an open system (mDICCO2-exch) is calculated
from the following mass-balance relationship:

( )
m
m m m m
D I C C O - e x c h2 =
⋅ − ⋅ − ⋅ −
− −−
δ δ δ
δ ε δ
1 3 1 3 1 3
1 3 1 3 1 3
2 3
C C C D I C
C C C
m e a s c a r b s o i l c a r b
s o i l C O C a C O c a r b
D I C D I C D I Cm e a s c a r b m e a s

In this equation, the value for mDICCO2-exch (moles of matrix-derived DIC that have exchanged with the
soil CO2) can be negative. This simply indicates that isotope exchange between DIC and the matrix is
the dominant exchange process.

The dilution factor then becomes:

q D I C D I C D I C
D I CF G
m e a s c a r b C O e x c h
m e a s
−
−=
− +m m m
m
2
Arranjo geológico do aquífero Bunter (arenito
Andrews et al., 1994).
arenitoarenito Bunter. A Bunter. A correlacorrelaççãoão inversainversa do do aa1414C C
com o com o δδ1313C C demonstrademonstra o o efeitoefeito das das reareaççõesões de de
nãonão decaimentodecaimento nana atenuaatenuaçção/diluião/diluiççãoão do do 1414CC
0
20
40
60
80
1 4 6 8 10 12 14 16 18 20 22 25 28 30
-14
-13
-12
-11
-10
-9
-8
a
14
C
D
IC
p
m
C
C δ 13C
‰
V
P
D
B
B
14C
13C
3H > 2 TU 3H - free```   