Analytical hydrogeochemistry
2. Stable isotopes in hydrogeology
Tento učební materiál vznikl v rámci projektu Rozvoj doktorského studia chemie
č. CZ.02.2.69/0.0/0.0/16_018/0002593
Outline
• Isotopes in hydrogeochemistry
– Structure of atoms
– Isotopes
• Radiogenic isotopes
• Cosmogenic isotopes
• Study methods
• Stable isotopes
– Equilibrium fractionation
– Fractionation coefficient
– Fractionation with respect to the standard
– Isotopic enrichment
– Kinetic fractionation
• Utilization of isotopes in hydrogeochemistry
Repetition
1. What is an isotope? Give an example.
2. What is isotope fractionation? Explain with
an example.
3. What physical properties are important in
isotope fractionation?
4. What does the value 𝛿2H= 2 ‰ tell me
about the heavy hydrogen content of 2H in
the sample?
STRUCTURE OF ATOMIC NUCLEI
Atomic model
• Core
– Protons and neutrons.
– Mass.
• Electron shells
– The electrons have a particlewave
character and move
with a certain probability in
the space defined by the
orbitals.
– Volume.
• Electroneutral particle
– Positively charged core.
– Negatively charged electrons.
Helium atom model
"HeliumatomQM"byUser:Yzmo-Ownwork.LicensedunderCCBY-SA3.0viaWikimediaCommons
-
https://commons.wikimedia.org/wiki/File:Helium_atom_QM.svg#/media/File:Helium_atom_QM.svg
Atom
• The properties of an atom are
determined by the number of
protons in the nucleus.
• Proton number Z.
• The proton number also identifies
the element to which the atom
belongs.
• Neutron number N.
• Total number of nucleons
(particles in the nucleus)
indicates the nucleon (mass)
number A, which determines
which isotope it is:
A = Z + N
• We then describe the element in
general
XA
Z
Helium atom model
"HeliumatomQM"byUser:Yzmo-Ownwork.LicensedunderCCBY-SA3.0viaWikimediaCommons
-
https://commons.wikimedia.org/wiki/File:Helium_atom_QM.svg#/media/File:Helium_atom_QM.svg
Nuclides and isotopes
• All atoms of one element
have the same proton
number, they can differ in
the number of neutrons.
• A nuclide is a substance
composed of atoms that
have the same proton
number and nucleon
number.
• Isotopes are nuclides that
have the same proton
number but different
nucleon numbers.
• The nuclide is generally
used in nuclear physics.
• An isotope is a narrower
term used in geochemistry
in the context of single
element nuclides.
UF,O,C, 238
92
19
9
18
8
12
6
OO,O, 18
8
17
8
16
8
Isobars and isotons
• Isobars
– Atoms with the same
number of particles in the
nucleus (p + n)
– E.g.: 40S, 40Cl, 40Ar, 40K, 40Ca
• Isotons
– Atoms with the same
number of neutrons in the
nucleus
– E.g.: 36S, 37Cl, 38Ar, 39K, 40Ca
Nuclear binding energy
proton: 1.007593 daltons = 1.6726231 × 10 –27 kg
neutron: 1.008982 daltons
electron: 0.000548756 daltons = 9.10093897 × 10 –31 kg
weight loss  = W - M
W - sum of masses of individual particles
M - mass of combined particles
4He = 2mp + 2mn + 2me = 4.034248 daltons
m(4He) = 4.003873 daltons
 = 0.030375 daltons (ie 28.28 MeV -> binding energy)
E = mc2
Nuclear binding energy
0
2
4
6
8
0 50 100 150 200 250
Hmotové číslo
Vazebnáen.jádra(MeV)
H
He
Li
B
Fe
U
Isotopes
Edited,originalByBenRG-Ownwork,PublicDomain,https://commons.wikimedia.org/w/index.php?curid=7900237
Isotopes
Taken from Clark (2015)
Isotopes
• There are at least 3,300 nuclides – naturally
occurring or experimentally characterized –
the vast majority of which are unstable and
rapidly decay. Only the stable and slowly
disintegrating isotopes can be found on Earth.
• There are two exceptions:
1. U and Th decay products,
2. radionuclides formed by cosmic radiation in the
atmosphere.
Isotopes
In terms of use in geology, we can divide
isotopes into three categories:
1. Stable isotope systems
2. Radiogenic isotope systems
– The decay of a radioactive isotope produces a
radiogenic isotope
3. Cosmogenic isotope systems
– Formed by bombarding the atmosphere with
cosmic rays (14C–14N)
Radiogenic isotopes
• They are formed by the decay of unstable nuclei
– E.g. 87Rb -> 87Sr
– End product or decay chain
• During decay, a large amount of energy is released
– Radioactivity
• The vast majority of existing isotopes are unstable
• Decisive role of half-life
• Use in geochronology, magmatic geochemistry, dating
of recent sediments, tritium
Cosmogenic isotopes
• Formed by cosmic rays
– 14N + n → 14C + p
• The radiocarbon method
– Quaternary geology and archeology
– The content of 14C in an organic sample relative to
the content in the atmosphere = age estimate
– Half-life 5730 years – suitable for dating up to
50000 BP
• Radiogenic 3He
PRINCIPLES OF ISOTOPE
FRACTIONATION
Introductory question
• The container contains liquid water and
gas in thermodynamic equilibrium.
• Most water molecules contain the
oxygen isotope 16O. Some contain the
isotope 18O. The ratio between the
amount of 16O and 18O in water is known.
• What will be the isotopic ratio between
16O and 18O in the gas?
a. The ratio of isotopes in the gas will be
the same as in the liquid.
b. Compared to the liquid, there will be
more oxygen 16O in the gas.
c. Compared to the liquid, there will be
more oxygen 18O in the gas.
Gas
Liquid
?
H − 18 O −H
H − 16 O −H
Liquid water in a closed container
Isotopic fractionation
• Enrichment of one phase by a given isotope
compared to another phase.
• The result is a different isotopic composition of
sea and rainwater, fossilized shells in different
parts of the formation, growth lines in cave
sinters, carbon dioxide in the atmosphere and
volcanic gases, etc.
• Several types of fractionation mechanisms
– Equilibrium fractionation
– Kinetic fractionation
– Rayleigh distillation
– …
Physical properties important for
fractionation
• Low isotope mass and relatively large mass
difference between isotopes
– The difference between 87Sr and 86Sr is 1%
– The difference between 18O and 16O is 12%
• High degree of covalent bonding – greater
effect of energy reduction.
• More oxidation states.
• Relatively high concentration of less
represented isotope (at least tenths of %).
Principle of mass fractionation
• If an object hangs on a
spring, it will oscillate
with higher frequency
than a heavier object.
H 2 O
Principle of fractionation
• When replacing a lighter
atom with a heavier atom,
the frequency (f) of the bond
vibrations decreases.
• The total bond energy
decreases according to
E = h × f
• The result will be lower
internal energy (U).
• The thermodynamic and
kinetic properties of the
molecule change.
Physical balance
• All molecules of an ideal
substance have the same
kinetic energy (EK).
EK = ½ m × v2
H − 18 O −H
H − 16 O −H
Physical balance
• All molecules of an ideal
substance have the same
kinetic energy (EK).
EK = ½ m × v2
• Heavier water molecules
have lower velocity and are
less readily released into gas
(and conversely, lighter
molecules more easily escape
from the liquid).
H − 18 O −H
H − 16 O −H
Physical balance
• All molecules of an ideal
substance have the same
kinetic energy (EK).
EK = ½ m × v2
• Heavier water molecules
have lower velocity and are
less readily released into gas
(and conversely, lighter
molecules more easily escape
from the liquid).
• A slight fractionation occurs –
the isotope ratio in gas and
water will be different.
H − 18 O −H
H − 16 O −H
Fractionation coefficient
• R a … ratio of the heavier isotope
to the lighter isotope in phase a
• R b … ratio of the heavier isotope
to the lighter isotope in phase b
α 𝑏
𝑎
=
𝑅 𝑎
𝑅 𝑏
gas
liquid
• Specific type of equilibrium constant
α 𝑔𝑎𝑠
𝑙𝑖𝑞𝑢𝑖𝑑
=
Τ18O 16O 𝑙𝑖𝑞𝑢𝑖𝑑
Τ18O 16O 𝑔𝑎𝑠
α 𝑔𝑎𝑠
𝑙𝑖𝑞𝑢𝑖𝑑
= 1.0092 (at 25 °C)
Difference with respect to the
standard
• Due to the extremely small variations in isotope
concentrations, the value is referenced with
respect to standards.
• How much the sample differs from the standard:
▪ 𝛿 < 0 … the sample is depleted of a heavier isotope
▪ 𝛿 > 0 … the sample is enriched with a heavier isotope
𝛿 =
𝑅 𝑠𝑎𝑚𝑝𝑙𝑒 − 𝑅 𝑠𝑡𝑎𝑛𝑑𝑎𝑟𝑑
𝑅 𝑠𝑡𝑎𝑛𝑑𝑎𝑟𝑑
× 103‰
Standards
element symbol ratio standard abs. ratio
H δD 2H/1H SMOW 1.557×10–4
If δ6 Li 6Li/7Li NBS L-SVEC 0.08306
B δ11B 11B/10B NBS 951 4.044
C δ13C 13C/12C PDB 1.122×10–2
N δ15 N 15N/14N ATM 3.613×10–3
O δ18O 18O/16O SMOW, PDB 2.0052×10–3
δ17O 17O/16O SMOW 3.76×10–4
S δ34S 34S/32S CDT 4.43×10–2
SMOW - Standard Mean Ocean Water (sometimes also VSMOW - Vienna Standard Mean Ocean Water )
PDB - Pee Dee Belemnite (C and O in carbonates, sometimes also VPDB)
ATM - ATMospheric nitrogen
CDT - Canyon Diablo troilite (from meteorite)
NBS - National Bureau of Standards (USA)
Example 1
• What is the value of δ18O for water vapor
above ocean surface (25 °C)?
𝛿18O 𝑣𝑎𝑝𝑜𝑟 =
Τ18O 16O 𝑣𝑎𝑝𝑜𝑟 − Τ18O 16O 𝑜𝑐𝑒𝑎𝑛
Τ18O 16O 𝑜𝑐𝑒𝑎𝑛
× 1000‰
vapor
water
α 𝑣𝑎𝑝𝑜𝑟
𝑜𝑐𝑒𝑎𝑛 =
Τ18O 16O 𝑜𝑐𝑒𝑎𝑛
Τ18O 16O 𝑣𝑎𝑝𝑜𝑟
α 𝑣𝑎𝑝𝑜𝑟
𝑤𝑎𝑡𝑒𝑟
= 1.0092 (at 25 °C) 𝛿18O 𝑜𝑐𝑒𝑎𝑛 = 0‰
The ocean is
the standard!
α 𝑣𝑎𝑝𝑜𝑟
𝑜𝑐𝑒𝑎𝑛 =
𝛿18O 𝑜𝑐𝑒𝑎𝑛 + 1000
𝛿18O 𝑣𝑎𝑝𝑜𝑟 + 1000
Example 1
• What is the value of δ18O for water vapor
above ocean surface (25°C)?
𝛿18O 𝑣𝑎𝑝𝑜𝑟 =
Τ18O 16O 𝑣𝑎𝑝𝑜𝑟 − Τ18O 16O 𝑜𝑐𝑒𝑎𝑛
Τ18O 16O 𝑜𝑐𝑒𝑎𝑛
× 1000‰
vapor
water
α 𝑣𝑎𝑝𝑜𝑟
𝑜𝑐𝑒𝑎𝑛 =
Τ18O 16O 𝑜𝑐𝑒𝑎𝑛
Τ18O 16O 𝑣𝑎𝑝𝑜𝑟
α 𝑣𝑎𝑝𝑜𝑟
𝑤𝑎𝑡𝑒𝑟
= 1.0092 (at 25°C) 𝛿18O 𝑜𝑐𝑒𝑎𝑛 = 0‰
The ocean is
the standard!
α 𝑣𝑎𝑝𝑜𝑟
𝑜𝑐𝑒𝑎𝑛 =
𝛿18O 𝑜𝑐𝑒𝑎𝑛 + 1000
𝛿18O 𝑣𝑎𝑝𝑜𝑟 + 1000
𝛿18O 𝑣𝑎𝑝𝑜𝑟 =
𝛿18O 𝑜𝑐𝑒𝑎𝑛 + 1000
α 𝑣𝑎𝑝𝑜𝑟
𝑜𝑐𝑒𝑎𝑛 − 1000‰ 𝜹 𝟏𝟖 𝐎 𝒗𝒂𝒑𝒐𝒓= −9.1‰
Enrichment factor
• Sometimes it is advantageous for us to express
the difference between the isotope content in
two substances
– E.g. mineral phases crystallized from the melt
• To compare isotope fractionation with δ–‰
values, we can express the fractionation factor
α as the enrichment factor ε
δ 𝑎 − δ 𝑏 = ∆ 𝑏
𝑎
𝜀 = 𝛼 − 1 × 103‰
Example
Source fractionation
α 𝑏
𝑎
=
𝑅 𝑎
𝑅 𝑏
α 𝑏
𝑎
− 1 =
𝑅 𝑎
𝑅 𝑏
− 1=
𝑅 𝑎 − 𝑅 𝑏
𝑅 𝑏
α 𝑏
𝑎
− 1 =
δ 𝑎 + 103 − δ 𝑏 + 103
δ 𝑏 + 103
=
δ 𝑎 − δ 𝑏
δ 𝑏 + 103
𝛿 =
𝑅 𝑎 − 𝑅 𝑏
𝑅 𝑏
× 103
𝑅 𝑎 =
𝛿 𝑎 𝑅 𝑏
103
+ 𝑅 𝑏 =
𝑅 𝑏 𝛿 𝑎 + 103
103
α 𝑏
𝑎
− 1 ≈
δ 𝑎 − δ 𝑏
103
if α 𝑏
𝑎
< 1.010 𝑡ℎ𝑎𝑛 δ 𝑏 + 103 ≈ 103
103 α 𝑏
𝑎
− 1 ≈ δ 𝑎 − δ 𝑏
δ 𝑎 − δ 𝑏 = ∆ 𝑏
𝑎
103 α 𝑏
𝑎
− 1 ≈ ∆ 𝑏
𝑎
Chemical equilibria
• Equations analogous to physical equilibria.
• The stronger covalent bond of heavier atoms results in
them preferentially entering compounds where they
are more strongly bound.
• As the temperature increases, the value of the
equilibrium constant decreases as the significance of
the difference in bond strength decreases.
• The difference in bond strength also affects the kinetics
of the reactions in which the bonds occur.
– Weaker bonds lead to a faster reaction of the lighter
isotope.
– Isotopic equilibrium also occurs in chemical equilibrium
over time.
Influence of temperature on
equilibrium fractionation
• Fractionation is a thermodynamic process – it depends on the temperature
• At high temperatures α is very close to 1
• High values for low temperature processes (surface conditions)
UsedfromClark(2015)
Geothermometry
TakenfromAllègre(2008)
At high temperatures, the fractionation is negligible.
Heavy oxygen δ18O
TakenfromAllègre(2008)
Heavy hydrogen δD
TakenfromAllègre(2008)
Kinetic fractionation
• Every process reaches equilibrium after a sufficient time.
• Equilibria disrupted by
– Chemical reactions
– Transport of substances
– Influence of temperature
– Biological processes
– And more…
• Out of balance, one direction of reaction significantly prevails.
• If equilibrium is not reached in the process, the process products are
enriched in a lighter isotope more than the equilibrium fractionation
coefficient would indicate.
• Each process reaches isotopic equilibrium after sufficient time.
• The question is how quickly the system reaches isotopic equilibrium.
• Isotopic exchange is given by:
– Temperature.
– It proceeds faster in liquids and gases (due to weak diffusion in solids).
– Position of a given isotope in the structures of compounds.
Kinetic fractionation
• Assume that the reaction A -> B
takes place and at the same time
fractionation takes place.
• If A and B are in contact long
enough, isotope exchange will
occur until the equilibrium is
finally reached.
• In order to maintain the kinetic
fractionation, the reactants and
products must not be in contact
(typically biota effects – removal
of reactants away from the cells).
• Temperature accelerates
fractionation – kinetic effects at
high temperatures are usually not
observable.
• The result is "isotopic memory" –
the solid phase can carry sealed
information about the kinetic
fractionation.
Taken from Allègre (2008)
Other effects on fractionation
• Diffusion fractionation
– The difference in velocity of heavier and lighter isotopes
– It dominates in hydrogeological insulators
• Microbially controlled processes
– Microorganisms prefer certain isotopes
– Lighter isotope bonds are more easily broken
Quantitative reactions
• Geochemical processes out of equilibrium
– Redox processes in groundwater
– Dissolution of minerals with the removal of products
– Ice melting
• The reaction affects the entire surface (structure)
– there is no preference for isotopes
Rayleigh fractionation
• Gradual fractionation of isotopes leads to ever
lighter/heavier products
• Evaporation of water from the tank – preferred
removal of light isotopes, water is getting heavier
𝑅𝑓 = 𝑅 𝑜 𝑓(𝛼−1)
• R f … Isotopic ratio of the reservoir after the reaction to
the residual fraction f
• R o … initial reservoir isotopic ratio (for f = 1)
• α … fractionation ratio
• Sometimes also Rayleigh distillation
Crystallization of magma
• A special case of a closed system – we assume that the crystals are
immediately isolated from the rest of the magma.
• Similarly, we can model water condensation and other systems.
fractionation coefficient
after integration
A and B are representative
isotopes
δ is the difference between the original isotopic composition and the
isotopic composition after it has condensed (crystallized) part f
Fractionation
Isotopic composition of the remaining vapor after condensation (f − 1) of the amount of water
compared to the original isotopic composition of the steam. Non-equilibrium and equilibrium
condensation are shown.
-50
-45
-40
-35
-30
-25
-20
-15
-10
-5
0
00,20,40,60,81
f (zbývající pára)
δ(‰)
α = 1,01
Rayleighova frakcionace
rovnovážná kondenzace
Temperature dependence
Dependence on composition and pressure: heavier isotope to a phase with a stronger bond (and covalent)
18O - more in quartz than in magnetite, 18O - more in CO3
2– than in water (30‰); pressure effect negligible:
(∂G/∂p) = ∆V
0,98
0,99
1,00
1,01
1,02
1,03
1,04
1,05
1,06
1,07
0 1 2 3 4
1000/T (K)
αCO2-H2O
α = 0,969 + 24,3/T
0
T (°C)
501002003005001000
0,99
1,00
1,01
1,02
1,03
1,04
1,05
1,06
1,07
0 5 10 15
106/T2 (K)
αCO2-H2O
α = 0,998 + 500/T2
0
T (°C)
501002003001000
The temperature dependence of the fractionation is usually determined experimentally for individual reactions
and usually takes the following forms:
UTILIZATION OF STABLE ISOTOPES
Application
• Monitoring of processes in the hydrological
cycle resulting in differences in isotopic
distribution:
– Evaporation and formation of atmospheric water
– Condensation and precipitation with decreasing
temperature
– Re-evaporation on continents
– Groundwater mixing
– Water-mineral-gas isotopic exchange (deep water,
long retention times)
δ18O and δD
• General observations
1. Fresh waters in warm areas are enriched and in cold areas
depleted of heavy isotopes
2. Strong linear correlation between δ18O and δD in rainwater –
slope value 8 and the intersection with the y-axis (D) 10 ‰
Taken from Clark (2015)
Precipitation fractionation
• Rayleigh fractionation
Taken from Clark (2015)
Precipitation fractionation
• Ideal development:
Taken from Clark (2015)
Correlation of δ18O and temperature in
precipitation
• Observable in longer-term comparison
(annual or monthly averages)
Taken from Clark (2015)
Correlation of δ18O and temperature in
precipitation
• Correlation indicates little development of
rainwater
TakenfromClark(2015)
Deviations from the ideal
• Problematic correlations of individual events
• Additional water supplied by evaporation
from rivers and lakes
– Mixing of water vapor in the troposphere
– Precipitation dynamics (vertical gradient of
condensation and transport)
• Relationships fit long-term averages well
• Description of empirical relationships
Meteoric water lines
• Global Meteoric Water Line (GMWL)
• Average of many local/regional MWL
– Various slopes and intersections
• Slope 8 = stronger fractionation of D
– Increases with lower temperature
Calculated curve Real data
• Observed MWL differs from calculations
• Consequence of non-ideal fractionation
Taken from Clark (2015)
Other influences on the MWL
• Mixing of water vapor of different composition as
it moves through the troposphere
• Evapotranspiration from areas of the continent
over which the water has passed
• Evaporation after condensation (especially if the
air is dry before precipitation)
– Another enrichment by 18O
– Necessary to take into account the precipitation
intensity – in dry regions significant shifts (extreme
effect of evaporation on rain) up to slope 7
– Weighted averages
Intersection with D
• MWL does not pass through the SMOW value, at the intersection the D
value is +10‰ (deuterium excess)
• It varies regionally in the range from ca. 0 to 20‰
• Influence of kinetic fractionation (nonequilibrium) – faster evaporation
than condensation (influence especially on 18O)
Taken from Clark (2015)
Differences in LMWL
• Complicates evaluation and comparison
Taken from Clark (2015)
Differences in LMWL
• Determining origin of water
Taken from Clark (2015)
Temperature effects in precipitation
• Distance from the ocean, altitude
• Distinctive origin of groundwater
TakenfromClark(2015)
Example
• Coast of Chile – scarce rainfall (100 mm/year)
• What can we say about the origin of waters?
TakenfromClark(2015)
Seasonal effects
• Altitude, distance from the sea and climate zone
• Differences in summer/winter rainfall composition
– observable in surface- and ground- water
Taken from Clark (2015)
Example
• Alaska – a small catchment area in the permafrost region
• What can we say about the origin of the water in the
stream?
Taken from Clark (2015)
Paleoclimatic applications
• Strong correlation of T and δ18O
• Isotopic composition of ice and snow
– Temperature changes over the last 650,000 years
• Values of δ18O in calcite precipitated from
groundwater (speleothemes)
• Groundwater from the end of the Pleistocene
trapped in the aquifers with deep circulation (the
high hydraulic gradient of the melting glacier
pushed them to a depth from which they wash
out extremely slowly).
Paleoclimatic applications
• Glacial waters in aquifers.
Taken from Clark (2015)
Groundwater recharge
• Infiltration-related processes dampen isotopic
variations in groundwater
– Water mixing, dispersion, evapotranspiration…
• Determination of the average isotopic
composition from the monthly averages is
problematic
• Weighted averages taking into account
precipitation totals
• We can usually simplify the composition of
groundwater to the annual weighted average
precipitation composition.
• In the tropics, the
seasonal precipitation
(monsoon) is strongly
distorted.
– The weighted average solves
this problem
• In the temperate zone,
groundwater is
replenished mainly by
melting snow and
precipitation in spring and
autumn.
– Due to evapotranspiration,
decreased summer
groundwater recharge
– The weighted average tends
to be heavier than
groundwater
• In arid areas, distortion
occurs due to the direct
evaporation of small
precipitation events.
– Solved by weighted average
TakenfromClark(2015)
Exercises
Plot the following points /
areas in the H2O isotopic
composition graph:
(a) Average seawater
composition (Standard
Mean Ocean Water -
SMOW)
b) Rainwater
c) Equatorial lakes and
rivers with high
evaporation
d) Water in polar glaciers
We can isotopically
distinguish three types
of water:
a) Marine
b) Precipitation
c) Geothermal
Geothermal waters
follow rainwater with
respect to δ2H, but
contact with silicate
rocks increases δ18O
Large amounts of water
interacting with the rock can
"imprint" their isotopic
composition (ie reveal the origin
of the fluid in the processes).
Taken from Gill (2015)
Taken from Gill (2015) by Craig ( 1961 , 1963) and Taylor (1974)
Hydrothermal systems
Fractionation of δD and δ18O
in meteoric hydrothermal systems.
Fractionation occurs due to
heating, boiling and water mixing.
-200
-180
-160
-140
-120
-100
-80
-60
-40
-20
0
-20 -10 0
δ18
O (‰)
δD(‰)
Yellowstone
Lassen Park
meteorická voda
chloridové
kyselé sulfidické
δ 13 C
The value of δ 13 C
Carbonates in sea water ~ 0 ‰
Atmospheric CO 2 −7 ‰
Organic matter Significantly depleted depending
on the mechanism of
photosynthesis (C3 and C4)
Plankton – tropics −20 ‰
Plankton – polar oceans −25 to −30 ‰
Terrestrial plants Very variable with a diameter of
around −25 ‰
Soil CO2 −20 ‰
CO2 in precipitation Negative, very variable
Carbonates formed in the oceans Close to zero
δ13C groundwater
• Determined by the balance of carbon sources and
sinks
• Main sources
– Dissolution of carbonate minerals (relatively heavy
carbon)
– Oxidation of organic matter (relatively light carbon)
– CO2 from the soil atmosphere (relatively light carbon)
• Metanogenetic redox processes
– Methanogenesis produces very light methane and
heavy CO2
– Oxidation of methane produces light CO2
Utilization of δ13C
• Carbon origin:
– Organic (light)
– Inorganic (heavy)
• Origin of secondary carbonate minerals
Radioactive isotopes
• Tritium, 14C
• Water dating – water with T will be younger
than 1950
• Radiocarbon method applicable to organic
substances in water
• Problem of water mixing!
18O in carbonates – the key to the
paleoclimate
• Calcite in equilibrium with
seawater is slightly
enriched by 18O.
• The fractionation
coefficient is strongly
temperature dependent.
• Oxygen ratios in marine
carbonates serve as
indicators of sea
temperature during
formation -> paleoclimatic
reconstruction.
α 𝑤𝑎𝑡𝑒𝑟
𝑐𝑎𝑙𝑐𝑖𝑡𝑒
=
Τ18O 16O 𝑐𝑎𝑙𝑐𝑖𝑡𝑒
Τ18O 16O 𝑤𝑎𝑡𝑒𝑟
α 𝑤𝑎𝑡𝑒𝑟
𝑐𝑎𝑙𝑐𝑖𝑡𝑒
= 1.0286 (25°C)
𝑇 ℃ = 16.5 − 4.3 × Δ 𝑤𝑎𝑡𝑒𝑟
𝑐𝑎𝑙𝑐𝑖𝑡𝑒
+ 0.13 × Δ 𝑤𝑎𝑡𝑒𝑟
𝑐𝑎𝑙𝑐𝑖𝑡𝑒2
Paleoclimate
• Urey et al. (1951):
fractionation of 18O between
calcite and water on the
example of Jurassic
belemnite from the Isle of
Skye.
• 4 summer and winter
seasons identified.
• Uncertainty is connected
with the composition of
seawater.
TakenfromGill(2015)accordingtoUreyetal.(1951)
Paleoclimate
δ18O in the sea is not constant, it changes
with temperature due to the deposition
of isotopically lighter water in glaciers
Glaciers of Greenland: −30 to −35‰
Antarctic glaciers: −50‰
The total variability of sea water is 1‰
Today's state:
Continental ice: 27.5 million km3
Water in the oceans: 1350 million km3
Ice ages: increase in ice by 42 million km3
level reduction by 125 m
pevnina
δ
18
O = – 30 ‰
interglaciál
oceán
δ18
O = 0 ‰
led
pevnina
δ18
O = – 30 ‰
glaciál
oceán
δ18
O = 1,5 ‰
led
δ18
O
0 1 2
Glaciers
• δ18O and δ2H in polar
ice changes depending
on the temperature
during snowfall.
• The Vostok well in East
Antarctica provides a
dated record for
420,000 years.
Fractionation in ice
Reconstruction of paleotemperatures from ice of Vostok well based on δD. The δ18O curve shows
changes in the isotopic composition of the ocean derived from sediment carbonates.
-10
-8
-6
-4
-2
0
2
0 50000 100000 150000
stáří (roky)
ΔT(°C)
-490
-480
-470
-460
-450
-440
-430
-420
-410
-400
0 50000 100000 150000
stáří (roky)
δD(‰)
-3,0
-2,5
-2,0
-1,5
-1,0
-0,5
0,0
0,5
1,0
δ18
O(‰)
Paleoclimatic data
CC BY-SA 3.0, https://commons.wikimedia.org/w/index.php?curid=184098
Example 3
• Analysis of Neogene sea limestone showed a
value of δ18O = −1.3 ± 0.1 ‰.
• Determine the temperature of the seawater in
which it was stored. We assume that the
isotopic composition of water was the same
as today.
δ 𝑎 − δ 𝑏 = ∆ 𝑏
𝑎
𝑇 ℃ = 16.5 − 4.3 × Δ 𝑤𝑎𝑡𝑒𝑟
𝑐𝑎𝑙𝑐𝑖𝑡𝑒
+ 0.13 × Δ 𝑤𝑎𝑡𝑒𝑟
𝑐𝑎𝑙𝑐𝑖𝑡𝑒2
Example 3
• Analysis of Neogene sea limestone showed a
value of δ18O = −1.3 ± 0.1 ‰.
• Determine the temperature of the seawater in
which it was stored. We assume that the
isotopic composition of water was the same
as today.
Example 3
• Analysis of Neogene sea limestone showed a
value of δ18O = −1.3 ± 0.1 ‰.
• Determine the temperature of the seawater in
which it was stored.
• How do the values change if we assume that
the variation in δ18O for water could be up to
1‰ (ie ± 0.5‰)?
Example 3
• Analysis of Neogene sea limestone showed a value of δ 18 O
= −1.3 ± 0.1 ‰.
• Determine the temperature of the seawater in which it was
stored.
• How do the values change if we assume that the variation
in δ18O for water could be up to 1‰ (ie ± 0.5‰)?
The resulting error is 22.3 ± 2.3°C
Stable carbon isotopes – the key to
history of life
• Seawater and carbonate
rocks are enriched in the
heavier isotope 13C (~
5‒10‰) compared to the
atmosphere.
• Organic matter contains
significantly less 13C -> δ13C
~−20‰.
• Preference of 12C in
photosynthesis by the
enzyme Rubisco.
• δ13C is an ideal tracker for
monitoring photosynthetic
processes.
Forms of carbon occurrence
• Inorganic carbon
– δ13C close to zero
• Reduced carbon formed by
the decomposition of
organic matter
– δ13C with significantly
negative values
Geochemical traces of the
development of life
2 disturbances :
1. The first
development of
photosynthetic
organisms associated
with the release of
oxygen into the
atmosphere = CO2
depletion.
Violation of the
greenhouse effect.
2. Transition of the first
photosynetizing
organisms to land =
acceleration of O2
increase.
Taken from Gill (2015) by Shields & Veizer (2002)
The origin of life
Bell et al. 2015 , available at http://www.pnas.org/content/112/47/14518.full.pdf
Archeology – nutrition and origin of people
• Trace elements and isotopes in
teeth reflect the isotopic
composition of food ingested
in childhood.
• This is a reflection of the
geology of the area.
• It allows to assess the origin of
archaeological finds of human
remains.
– Migration and cultural
interactions.
• Nitrogen and carbon isotopes
can reveal dietary details.
– C3 plants: rice, fruits, tubers,
nuts, vegetables
– C4 plants: millet, sugar cane
Tollense Battlefield – dated by 14C to 1250 BC. Evidence of a battle of
surprising volume (up to 4000 participants) for the time. Stable isotopes
prove their very varied origin.
Archeology - sources of materials
• Origin of stone
artifacts (corneas,
obsidians, etc.)
• Origin of metals,
glasses or pigments.
• Reconstruction of
trade routes.
Archeology - sources of materials
Albarède et al. (201 6) Geochem. Persp. Flight. 2 , 127-137 | doi: 10.7185 / geochemlet.1 613
Silver isotope compositions of Roman silver coins pre- and post-dating the 211 BC monetary reform. Solid symbols:
denarii ( top ) and quadrigati ( bottom ). Open symbols: victoriati . Mint ages and uncertainties are from Crawford
(1974). Error bars on silver isotope proportions are the same as or smaller than the symbol size. Events listed on the
right-hand side are chosen for historical relevance. See Appendix for the range of Spanish values.
Tento učební materiál vznikl v rámci projektu Rozvoj doktorského studia
chemie
č. CZ.02.2.69/0.0/0.0/16_018/0002593
References
• Unmarked images are author's or taken over with the
permission of doc. Zeman and doc. Faimon
• Allègre , CJ (2008). Isotope Geology. Cambridge University
Press . ISBN: 9780511451126.
• Clark , I. (2015). Groundwater Geochemistry and Isotopes .
CRC Press . 442p. ISBN 978-1-4665-9174-5 ( eBook - PDF)
• Gill, R. (2015). Chemical Fundamentals of Geology and
Environmental Geoscience . 3rd Edition . John Wiley and
Sons . 288p. ISBN: 978-0-470-65665-5