stable isotopes & paleoclimates · pdf fileoxygen isotopes : seasonal variations in modern...
TRANSCRIPT
1
0
1
2
3
4W
ater
dep
th (
Kilo
met
ers)
Reconstructing Earth Historywith Stable Isotopes & trace
elements of Marine MicrofossilsForaminifera tests~ calcite (CaCOCO3)
MassSpectrometers
CO2
Isotopes:13C/12C 18O/16ONotation: δ13C (‰) δ18O (‰)
∆∆ Ocean Temperature Ocean TemperatureAs T increases, δ18Odecreases 1‰ ~ 4°C
Basic Temperature EquationT=16.9-4.0 (δ18Oc- δ18Osw)(Shackleton, 1974)
δ18O (‰)-3 -2 -1 0 1 2 3
0
1
2
3
4W
ater
dep
th (
Kilo
met
ers)
Mixed layer planktonicforaminifera
Benthic foraminifera
25 20 15 10 5 0
Temperature (°C)
Stable Isotopes & Paleoclimates
• O & C Basics– Notation– Fractionation
• Greenhouse Case Study• Icehouse Case Study
Who is this person?Dr. Harold Urey
2
What’s an Isotope?• Isotopes - of an element - same number of protons (P),
different number of neutrons (N)– (from the Greek isos, meaning “same,” and topos, signifying
“place”)
nmEm = mass number (A)n = atomic number (protons)A = P + N
16O, 17O, 18OP/N : 8/8, 8/9, 8/10
12C, 13C P/N : 6/6, 6/7
•Extranuclear structure remains the same!
What is Isotopic fractionation?
Knowledge of isotopic fractionation is derivedmainly from?
• laboratory calibration studies:– equilibrium precipitation experiments– exchange or kinetic effects between different
phases or speciessupplemented with:
• semi-empirical calculations - data and statisticalmechanics
• measurement of natural samples whoseformation conditions are well known
partial separation of isotopes between two substancesduring physical and chemical processes
3
Notation for Expressing Isotopic Compositions(abundances)
Delta Value (δ) - differences in ratios• Geochemical applications - difference in
absolute isotopic ratios betweensubstances is adequate
– more accurately measured thanabsolute ratios
• Ratios reported relative to a standard (ref.gas) in the delta notation expressed asparts per thousand or permil (‰)
10.0‰, = 1.0%R = absolute isotope ratios (heavy/light)13C/12C18O/16O
δA (‰)= RAsamp − RSTD
RSTD
x 103
δA(‰)= ( RAsamp
RSTD
-1) x 103
δ13C(‰)= 13C/12C(samp)!-!13C/12C(std)!
13C/12C(std) x 103
δ13C(‰)= ( 13C/12C(samp)!
13C/12C(std) -1 ) x 103
If Ra > Rb, heavier, higher, more positiveδ18O, δ2H or δD, δ15N, δ34S,
Why not absolute values?• convenience: working w/ small differences
(i.e. 13C = 1.1230% 1.1210%)• ensures greater consistency over time and between laboratories
-differences can be measured with greater precision
Equilibrium FractionationDefined as a redistribution of isotopes of an element among
various species or compounds– aA1 + bB2 ⇔ aA2 + bB1– where A & B represent two species, subscripts 1 &2 isotopes
that are substituted, a & b # of molesFor example; 1/3 CaC16O3 + H2
18O ⇔ 1/3CaC18O3 + H216O
Requirements for Equilibrium?• closed system and/or well mixed• Stable (on some time-scale that is rate dependent)• complete equilibrium is achieved when forward / backward reactions
are equal
• Highly temperature dependent !
4
General observations of Equilibrium Fractionation:
• Oxidation State:– heavier isotope accumulates in species/compound with higher
oxidation statei.e., SO4
-2 (S+6)is enriched in 34S relative to SO3 -2 (S+4) and FeS (S-2)
• Multiple phases:– higher density accumulates heavier isotope– δ18OS > δ18Ol > δ18Og
Condensation - heavier isotope is concentrated in the liquidphase, lighter isotope in the vapor phase
– For both δ13C and δ18O ~ CO2 < HCO3 < CaCO3• Temperature:
– Increase T difference in equilibrium isotope compositionsdecreases.
• Fractionation factor (α ) - defines the fractionation associated w/equilibrium exchange reaction between two substances.R = N*/N, N*=heavy isotope, 18O/16O
αA-B = RA
RB
1/3 CaC16O3 + H218O ⇔ 1/3CaC18O3 + H2
16Oα CaCO3 -H2O = 1.031 at 25°C
General observations concerning α• can be determined experimentally and theoretically• values tend to be close to unity (~1.00xx)• sign and magnitude dependent on many factors
• temperature (most important)• chemical composition• chemical structure• pressure
5
• Conversion of α to delta:δA= RA
RSTD
− 1
x 103 (‰)
and
δB= RB
RSTD
− 1
x 103 (‰)
αA-B = Ra
Rb
αA-B = δ A + 1000δ B + 1000
Express Ra & Rb in terms of δ and substitute into;
If the δ18Ol of the ocean is 1.0‰, and the water temperature is 25°C,assuming equilibrium, what is the δ18Oc of calcite that precipitates fromthis water? α c-l = 1.028 at 25°C
δ18Oc = α(1.0‰+103) - 103
δ18Oc = 29.03‰
Oxygen Isotopes-Temperature and Ice-volume
• most abundant element on earth(gas, liquid, solids)
• three stable isotopes16O - 99.763%17O - 0.037518O - 0.1995
• omnipresent, involved in mostnaturally occurring geologicprocesses
- sea water- hydrologic cycle- biosphere
• geothermometer– Partitioning of isotopes is
temperature sensitive
6
Isotope Ratio Measurements1. Gas Source Mass Spectrometers2. Convert Sample to CO2
– Carbonates -• reaction with 100% phosphoric acid (H3PO4) at
temperature to create CO2
CaCO3 + H3PO4 ⇒ CO2 +H2O + CaHPO4
• reactions temperatures - 25-90°C• advantages of higher temperature:
- increase reaction rates- decreases fractionation difference between carbonates
• O-isotope fractionation during reactionα CO2-CaCO3 = 1.01025 at 25°C
2. Waterequilibration w/ CO2 at a fixed temperature (25°C)
– Must know the fractionation factor:αCO2 -H2O = 1.041 at 25°C
• H2O >> CO2
Standards (18O/16O):
Standard abs. ratio(x106)
δ18O
v- SMOW Standard Mean Ocean
Water
2005 (±0.43)373
0‰
v SLAP Stand. Light Ant. Precip -55.5‰
v PDB Pee Dee Belemnite 2067 (±2.1) 28.64 (-2.20PDB)
Conversion - SMOW and PDB scalesδSMOW = 1.03086 δPDB + 30.86δPDB = 0.97002 δSMOW - 29.98
7
O-Fractionation Mechanisms• phase transitions of water - most
effective means of fract.– vapor - liquid - ice
• evaporation• condensation
– Vapor Pressure and, to a lesserextent, Freezing PointDifferences: Mass dependant
– H218O (20) < H2O (18)
• For liquid/vapor: O isotopefractionation decreases withincreasing temperatureFor 0-350°C (Horita & Wesolowski,
1994)1000 ln α l-v = -7.685+6.7123(103/T) -
1.6664 (106/T2) + 0.35041(109/T3)
5
6
7
8
9
10
11
12
0 20 40 60 80 100 120
18O/16O Fract. (liquid-vapor)
1000ln a (F&O,77)1000ln a (H&W,95)
1000
ln α
l-v
°C
8
* O always enriched in the C bearing species.CO2 dissolution (hydrolysis):
CO2 (g) ⇔ CO2 (d) + H2O ⇔ H2CO3 ⇔ H+ + HCO3 ⇔ H+ + CO3
•Assuming each species is in isotopic equilibrium with water a species (19°C) 103ln αa-H2OH2CO3 38.7HCO3 34.5CO3 18.2CO2 (g) 41.6
Partitioning of O between dissolved C & H2O
Paleothermometry in Marine Carbonates
• Isotope Exchange reaction between H2O andHCO3, CaCO3 enriches calcite in 18O.
Ca + HCO3 ⇒ CaCO3 + H2O + CO2
CaCO163 + H2O18 ⇒ CaCO18
3 + H2O16
αCO3−H2O = Rc/Rw = 1.028 at 25°C– Foraminifers are roughly 28‰ enriched relative
to seawater-– Fractionation can deviate for several reasons!
• Vital effects - inclusion of metabolic CO2
• Precipitation rates - move away fromequilibrium
– First temperature equation (Epstein et al. 1953)• T°C=16.5-4.3(δcc- δw) + 0.14(δcc- δw)2
δcc== CO2 at 25°C
9
Temperature dependence - fractionation increases with coolertemperatures
1000 ln αc-w = 2.78!x!106!!T2 - 2.89 (3)
(Erez and Luz, 1983).
The fractionation factor α c-w(Freidman & O’Neil, 1977)
Oxygen Isotopes : Seasonal Variations in Modern Foraminifera
• Plankton tows, Bermuda (Williams et al. , 1981, Palaeo3, v.33, p. 71-102)
10
Oxygen Isotopes: Lab and Field studies of Modern Foraminifera
• Lab culturing experiments ∆Orb-H2O = 3.19-0.208T
•Plankton Tows & Box Cores, Indian Ocean (20°N-30°S)∆Orb-H2O = 3.50-0.214T
(Bouvier-Soumagnac et al., 1985, JFR, v.15, p. 302-320)
Oxygen Isotopes: Pelagic Benthic Foraminifera
11
Oxygen Isotopes : Modern Benthic Foraminifera
•Box Cores•Rio Grande Rise, Norwegian Greenland Sea, E. PacificRise(Belanger et al.,1981, Palaeo3, v.33, p.205-220)
C. wullerstorfi Oridorsalis Pyrgo
Oxygen Isotopes: Modern Benthics
• Box Cores• California margin(Grossman, 1984, Palaeo3, v. 47, p.301-327)
12
Heterococcoliths
Calcidiscus leptoporus Emiliania Huxleyi
Coccoliths: Oxygen Isotopes• Culturing studies
– multiple species– 6‰ range
• Calciteδ18O variations as afunction of:– Growth rate– Temperature
• Species dependant
Ziveri, Stoll et al., 2003, EPSL, v. 210, p.137
13
Coccoliths: temperature
Ziveri, Stoll et al., 2003, EPSL, v. 210, p.137
Natural Abundances of Oxygen Isotopes: Water & Carbonates
14
O-isotope of precipitation• For liquid/vapor: O isotope fractionation increases with decreasing
temperature– for 0-100°C (Friedman & Oneil, 1997):
• Condensationδ18Ol > δ18Ov
Clouds become progressivelydepleted in 18O
δ18Ov decreases toward highlatitudes
Ice-sheet δ18O = -30 to -50
O-isotopes Natural Abundances: General Observations• Ocean Water
– Open Ocean - Relatively homogeneousδ18O = 1.5 to - 0.5‰δD =12 to -8‰
evaporation/precipitation/runoffδD = M δ18O
where M is the slope of linear trajectories (decreases w/increasing evaporation/precipitation)
• M ~ 6 (low latitudes) to 8 (high latitudes)– equatorial - P > E– subtropical - E > P– high lat - P > E
– N. Atlantic - end member -21‰• Deep Ocean - δ18O - mixing line between 0 and -0.5‰• Marginal Seas − δ18O - -3.0‰
high freshwater input
15
The “Salinity” Effect and δw
∆ Evaporation/Precipitation Rates
δ18O of precipitation
Condensation/Precipitationδ18Ol > δ18Ov
Clouds become progressivelydepleted in 18O
δ18Ov decreases toward highlatitudes
Ice-sheet δ18O = -30 to -50
16
Glacial Ice-volume and Ocean δ18O
Ice-sheets (18 kya, present day)
17
Sea-level Change, 150 kya to present
120meters
18
Carbon IsotopesC12 - 98.89%C13 - 1.11%
δ13C(‰)=13 C 12 C samp
13 C 12 C sta
− 1
x 103
Atlantic δ13CGlacial to interglacial cycles
Fractionation Mechanisms
1. Partitioning of C between CO2- and other hydrated speciesCO2 (g) ⇔ CO2 (d) + H2O ⇔ H2CO3 ⇔ H+ + HCO3 ⇔ H+ + CO3
– temperature dependent (~small):1000 ln a = A + B(106/T2)– When in isotopic equilibrium with CO2 (g)
Species A B 103ln α= ∆H2CO3- CO2 (g) -0.91 0.0063 -0.8HCO3- CO2 (g) -4.54 1.099 7.7CO3- CO2 (g) -3.4 0.87 6.4CaCO3- CO2 (g) -3.63 1.194 9.8
Deines et al. (1974)
CO3 can combine with divalent cations to form various carbonate minerals 2 sources for disequilibrium:
• precipitation rate - as rates increase the degree of fract. decreases• vital effects - metabolic C from respired CO2 may be incorporated
19
C fractionation w/ respect to HCO3
20
Carbon Isotopes: Natural Abundances
depletion of 13C in lipids relative to biomass as a function
Kinetic Effects: Photosynthesis∆CH2O-CO2 ~ -10-25‰ (all values are in terms of discrimination against 13C)• Why is plant δ13C lower than that of CO2 in the atmosphere?• Two key steps where fractionation can occur:
CO2(a) ⇔ CO2(i) ⇒ R-COOH1. diffusion - 12CO2 diffuses more rapidly2. bonding strengths - 12CO2 weaker (lower activation energy)
Uptake and conversion of CO2 involves three steps1. Diffusion - CO2 is transported through a boundary layer and the stomata into
a the internal gas space• ∆a-I = δa − δi = 4.4‰
2. Dissolution - CO2 dissolves into cell sap and diffuses to chloroplast (as HCO3in C4)• isotope fractionation ~-0.7-1.1‰ (i.e, 13CO2 is less soluble)
3. Carboxylation - carbon fixation• chloroplast - enzymes - catalysts for fixing CO2• isotope fractionation variable:
C3 vs. C4
21
22
Planktonic Foraminifera Isotopes as a function of size
symbiotic and non-symbiotic plankton foraminifera (Orbulina universa and Globigerina bulloides(Spero et al., 1997, Nature)
Modern Benthic Foraminifera δ13C
Grossman, 1984, GCA, v.48, p. 1505-1512
DIC δ13C
Porewater δ13C
23
Modern Benthic Foraminifera δ13C
Belanger et al. 1981, PPP, v. 33, p.205-220.
Heterococcoliths
Calcidiscus leptoporus Emiliania Huxleyi
24
Coccoliths: Carbon Isotopes• Culturing studies,
– Constant T (17°) & light– multiple species– 5‰ range
• Calcite δ13C variations as afunction of:– Growth rate– Cell diameter– Surface area/volume
Ziveri, Stoll et al., 2003, EPSL, v. 210, p.137
Coccoliths: Sources of IsotopeDisequilibrium
Three potential effects:(1) kinetic fractionation during calcite precipitation or
cellular uptake(2) fractionation of the intracellular C pool during
photosynthesis(3) pH-dependent fractionation