soil isotopes final

Applications of stable isotopes in paleosols

- PaleoCO2 concentrations (dC13)

- Paleoecology (dC13)

- Paleoclimate (dO18)

- Paleoelevation (dO18)

- Paleotemperature and aridity (clumped delta47 and O17)

What material in soils do we analyze for their stable isotopic composition?

- soil organic matter

- authigenic minerals (minerals that form in soils under environmental conditions - soil carbonate, clay mienrals, siderite)

Father of Isotope Chemistry

Harold Urey

Father of Mass Spectrometry

Alfred Nier

Isotope

atoms with the same number of protons but different number of neutrons, giving them a different atomic mass

Stable isotope

an isotope that is stable and will not decay

Isotopologue

isotopically substituted molecule (ex. C(12)O2, C(13)O2, C(14)O2)

light environmental stable isotopes

large mass difference of light stable isotopes influences their velocity, and therefore their rates of physical and chemical reactions and diffusion

heavy stable isotopes

- mass difference between heavy isotopes is small (<2%)

- all isotopes behave similarly during physical and chemical reactions (important as tracers and for radiometric dating)

characteristics of stable light isotopes

- elements have relatively low atomic mass (atomic number <20)

- large relative mass difference between the rare (heavy) and abundant (light) isotope

- element form chemical bonds that have a high degree of covalent bonds

- elements generally exist in more than one oxidation state and form a wide range of compounds

- abundance of element and rare isotope are sufficiently high (a few tenths of a percent) to ensure precise analysis on a mass spectrometer

main applications of light stable isotope geochemistry

- thermometry

- tracers

- reaction mechanics

- paleoclimatology

fractionation

the partitioning of molecules with different masses due to a disparity in their relative reaction rates

fractionation types

1) physiochemical (equilibrium and non-equilibrium/kinetic process)

2) diffusive (non-equilibirum/kinetic process)

physiochemical fractionation

- phase change or chemical reaction

- basis of physiochemical fractionation is the difference in the strength of bonds formed by light and heavier isotopes of a given element

- greater dissociation energy is needed for heavier isotopes, which have stronger bonds; lighter nuclei react more quickly

diffusive fractionation

- diffusion of atoms across a concentration gradient

- fractionation arises from differences in the diffusive velocities between isotopes (ie lighter isotopes diffuse faster)

- possible to establish a steady state diffusion regim, but NOT an equilibrium (rate of forward reactions cannot = backward reactions)

- diffusion is a kinetic fractionation

temperature effect on fractionation

- greater fractionation at low temperatures

kinetic (non-equilibrium) fractionation

- occurs when thermodynamic equilbrium does not exist, usually as a result of fast, incomplete, or unidirectional processes (evaporation, diffusion, dissociation reactions, biologic reactions)

- effects will vary depending on reaction pathway, but typically enhance fractionation

stable isotope notation

d - difference of a sample relative to a standard (not a ratio)

reference standards

- VSMOW (Vienna - Standard Mean Oceanic Water)

- VPDB (Vienna Pee Dee Belemnite)

VSMOW

- used for all dO18 and dD analyses of water

- sometimes used for dO18 of calcite (only when comparing records with water or ice data)

VPDB

- calcite precipitated from ocean water

- used for most dO18 analyses of calcite and all dC13 analyses

- generally water is not converted to PDB scale for comparison

fractionation factors

a

- refers to the amount of fractionation (separation based on mass differences) that occurs during a specific physiochemical reaction at a given temperature

fractionation factor a>1

heavier isotopes incorporated in product

fractionation factor a<1

heavier isotopes retained in reactant

equilibrium isotope fractionation

- eq fractionation between two phases generally decreases with increasing temperature

- degree of fractionation is generally larger for elements whose mass difference is large

- heavy isotope is preferentially partitioned into site with stiffest bonds (strong and short)

- bond stiffness increases qualitatively with: high oxidation state, lighter elements, covalent bonds, a low coordination number

controls on dC13 values of soil carbonates

- Fractionation of atmospheric CO2 by vegetation type (grasslands - more positive, trees/shrubs - more negative)

- respiration rates - mixing of atmospheric CO2

- atmospheric pCO2 and dC13

- temperature

stable isotopic composition of soil carbonate

- C13 is preferentially incorporated into calcite during precipitation

fractionation of atmospheric CO2 by vegetation

- kinetic diffusion in soils

- C12 atoms have higher kinetic energy and diffuse faster than C13 atoms

- C(12)O2 molecules diffuse faster in a soil environment

- causes enrichment of about 4.4 per mil between soil respired CO2 and soil CO2

dC13 of C3 vegetation

-27.1 per mil

dC13 of C4 vegetation

-12.1 per mil

dominant C3 vegetation zones

- polar deserts, tundra, conifer forests, tropical/temperate broad-leafed forests

mixed C3/C4 vegetation zones

tropical temperate deserts, semi-desert

dominantly C4

need water stressed environment during growing season --> tropical temperate grasslands

- not all deserts have C4 (Mojave and Mediterranean)

soil carbonate values

C3: -12 to -10 per mil

C4: 0 to 2 per mil

soil respiration

release of CO2 from soil surface into atmosphere, caused by decomposition of organic matter, plant root respiration, and soil fauna respiration

respiration rates effect

- high soil respiration = more negative value

- low soil respiration = more positive value

atmospheric CO2 concentration effect

- low CO2 = more negative value

- high CO2 = more positive value

problems with soil carbonate in paleosols

- identifying soil carbonate

- diagenetic alteration of dC13 values

- estimating dC13 of atmosphere

- assumptions of soil respiration rates and temperature

controls on the dO18 of soil authigenic minerals

- dO18 of precipitation

- evaporitive enrichment of soil water

- temperature

- diagenetic alteration

stable isotopes and the hydrologic cycle

- lighter molecules evaporate more easily and diffuse to the surface faster than heavier molecules

- heavier molecules will precipitate more readily

dO18 of meteoric water: source

- dO18 of water body (ocean or lakes)

- temperature and humidity during evaporation

dO18 of meteoric water: sink

- dO18 of air mass evolves (becomes more negative) as it moves away from source

- temperature and humidity (seasonality)

- rayleigh distillation of air mass (continental effect, altitude effect, amount effect, seasonality, temperature)

rayleigh distillation

- isotopic composition of a reservoir changes as material is progressively removed in a fractionating process

- heavier isotope is preferentially entering liquid phase and progressively removed from clouds

altitude effect

- dO18 value of precipitation gets more negative with elevation

- function of rayleigh distillation and cooling of air mass

- avg lapse rate of -3 per mil/km

amount effect

- dO18 value of precip becomes more negative during higher rainfall intensity

- rayleigh distillation, change of water vapor isotopic value below clouds, no evaporative enrichment of rain droplets (rainfall can be modified during precip)

- cause of seasonal variation over island stations

continental effect

- isotopic values of precip become more negative toward the interior of a continent

- mostly same processes as altitude effect (rayleigh distillation and temperature)

seasonality

- large differences in seasonal temperature at mid-latitudes influence rayleigh distillation

- different source regions of moisture

- differences in evapotranspiration flux over continents

atmospheric temperature

- warmer temperatures = more positive dO18 of precip

- cooler temperatures = more negative dO18 of precip

evaporative enrichment

- dO18 of soil carbonate is more positive in more arid environments

- more evaporation of lighter isotopes = more positive values

soil temperature

i lowkey do not know

biggest assumption when using oxygen isotopes to infer past climatic or environmental change

assuming that precipitation doesn't change

why not use conventional dO18 paleothermometry?

- requires isotopic composition of water from which the carbonate grew (temp. estimates only as accurate as assumptions of isotopic comps. of past waters)

- diagenesis: carbonate minerals commonly undergo post-depositional transformations

isotope rules

- lighter molecules vibrate faster

- lighter molecules form weaker bonds

- isotope effects are stronger for bigger relative mass differences

- isotope effects are stronger at lower temperatures

clumped isotopes - isotopologues

- molecules that are identical except for their isotopic composition

- clumped isotopologues have multiple heavy isotope substitutions

- clumped arrangements are energetically favored, with more clumping at lower temps

clumped isotopes

isotopologues that include multiple rare, heavy isotopes

ex. C(13)O(18)O(16)

where C13 and O18 are the rare, heavy isotopes

challenges with clumped isotopes

- sample preparation (need large samples, slow, low abundance of target isotopologues, need to not scramble isotopologues)

- instrumentation (need high resolving power, need high resolution to measure small isotopic signals, need high abundance sensitivity)

- need at least 3 replicate samples to reduce uncertainty

clumped isotope applications

- paleotemperature estimates where source isotopic composition is unknown

- stability of ocean dO18 through time

- paleoelevation

- body temperature of extinct animals

- not for speleothems or fast-growing corals, mollusks unsure

clumped isotopes in formations

- marine carbonates: mostly form at equilibrium temperatures

- lake carbonates and soils: form at equilibrium temperatures but strongly biased to summer months

- freshwater shell: idk

- speleothems and travertines: do not form at equilibrium temperatures due to kinetic effects

diagenesis of clumped isotopes

- clumped isotopes are thermally reset below 2-3km (100-125C) burial

- if widespread, limits techniques to young, shallowly buried stratigraphic sections

dO18 of meteoric water can track

- moisture source

- how much it rains (amount effect)

- height of ancient mountains

- does NOT take into account evaporation

when we have dO18 of soil carbonate and want dO18 of meteoric water, we assume:

1) growth temperature

2) evaporation

triple oxygen isotopes

- tool to detect evaporation in waters and carbonates

- ambiguities in dO18 can be constrained with deltaO17

- can constrain hydrologic processes in geologic time

humidity related to deltaO17

- humid environments will be more positive per meg (ex. Huron River, MI is +15 per meg)

- arid environments will be more negative per meg (ex. Great Salt Lake, UT is -12 per meg)

O17 will be most useful in identifying:

1) when variation in dO18 of soil carbonate is due to increased evaporation

2) when variation in dO18 of soil carbonate is due to a difference in recharge dO18

kinetic diffusion of soil carbonate

add 4.4 per mil to account for kinetic diffusion of carbon isotopes in soil carbonate

temperature lapse rates for dC13 and dO18

dC13: 0.1 per mil/1C

dO18: 0.22 per mil/1C

what creates high isotopic values for soil carbonate?

- low soil respiration rates

- much colder values

- arid environments --> lighter isotopes evaporate faster

- shallow carbonate formation