EOS 335 Part II Flashcards
elements with at least 1 stable isotope
80
known stable isotopes
250
element with most stable isotopes
Tin, 10, 112Sn - 124Sn
elements with 8 stable isotopes
none
elements with 9 stable isotopes
only xenon
mononuclidic elements
27 (single isotope)
elements with at least 1 stable isotope
H - Pb (1-82) except technetium and promethium
elements without stable isotopes
> 82
stables isotopes are what state
ground state of a nuclei
isotopic elements of geochemical/biological interest have
2+ stable isotopes
lightest generally in greater abundance
(C,H,O,N,S)
isotope properties
same protons and electrons = same chemical behaviour
physiochemical differences
isotope properties, same chemical behaviour
enter same chemical reactions
form same bonds
rare isotope can trace abundant isotope
lighter stable isotope
generally more abundant
not Li, B
Important stable isotopes
H, D, C, N, O, S, Cl
isotope properties, physiochemical differences
lead to differences in distribution between phases boiling pt freezing pt density vapor pressure, etc.
Characteristics of elements for isotope effects
relatively low atomic mass
relative mass difference between rare and abundant is large
form chemical bonds w/ high degree of covalent character
abundance of rare is sufficiently high
can exist in more than one oxidation state
low atomic mass, isotope effects
H, He, C, N, O, S, Cl
exception - Fe isotopes fractionated by bacteria
large mass difference, isotope effects
∆m D-H = ca. 100%
∆m 13C-12C = 8.3%
∆m 18O-16O = 12.5%
why measure stable isotopes as ratios
utility- compare identical species/phases
measurement - measuring ratios increases precision
δ
‰
unitless
differences between sample and standard readings
not absolute isotope abundance
natural abundance standard
defined as δ = 0
Stable isotope ratio notations
δ^n X sample (‰) = [Rsample - Rstandard/ Standard] x1000
R
absolute abundance ratio
atom% ^n X / atom% ^m X
e.g. atom% 15N /atom % 14N
light/heavy wording
lots of heavy isotope = enriched, heavy less of heavy isotope = depleted, light e.g. more 13C = heavy, enriched, + less 13C = light, depleted, - 0 = standard
hydrogen stable isotope standards
SMOW, VSMOW
standard mean ocean water
oxygen stable isotope standards
SMOW VSMOW
standard mean ocean water
carbon stable isotope standards
PDB VPDB
Pee Dee Belemnite
13C/12C = 0.0112372
18O/16O = 0.0020671
nitrogen stable isotope standards
atm N2
atmospheric nitrogen
sulphur stable isotope ratio
CDT
Canyon diablo triolite
δ13C atmospheric CO2
-8.2‰
δ13C plants, kerogen, coal
-8 - -55‰
𝛿 13C oil
-22 - -50‰
𝛿13C natural gas
-25 - -100‰
range in hydrogen isotopic variation
-700 - +200‰
range in carbon isotopic variation
-140 - +40‰
range in nitrogen isotopic variation
-60 - +50‰
range in oxygen isotopic variation
-30 - +30‰
range in sulphur isotopic variation
-50 - +40‰
Hydrogen isotope mass differences
99.8%
Carbon isotope mass differences
8.36%
nitrogen isotope mass differences
7.12%
isotopes are subjected to
Isotope effects
higher mass differences =
larger isotope effects
more strongly fractionated
Isotope effects
departure in isotope ratio from global average abundance due to physiochemical mechanisms
types of isotope effects
EIE - equilibrium isotope effects
KIE - kinetic isotope effects
Isotope fractionation
expression of isotope effects
Oxygen mass differences
∆m 16O/18O = 12.%
sulphur isotope mass differences
∆m 32S/34S = 6.24%
causes for isotope effects
chemical, physical properties of isotope
physiochemical properties of isotopes
isotopologues
chemical, physical properties of isotopes
arise from differences in atomic mass
reaction rates
diffusion rates
equilibrium constants
physiochemical properties of isotopes
result of quantum mechanical effects
energy of molecule restricted to discrete E levels
e.g. heat capacity, density, vapour pressure
isotopologues of isotopes
same molecules with different masses
different vibrational energies
e.g. H2, DH, D2
different isotopologues
different masses = different vibrational energies = different zero-point energies
ZPE
zero point energy
energy difference between minimum in potential energy curve and ground state energy (Eo)
Eo =
1/2 hv
h = plancks constant
v = vibrational frequency
ZPE, heavy isotopologue
lower ZPE than lighter isotopologue because v varies inversely with mass
isotopologue bonds
weaker in light isotopologues, easier to break, due to higher ZPE
difference in chemical properties of isotopologues can be evaluated in terms of
vibrational frequencies
transition states, in case of kinetic effects
energy well, isotopologues
heavier isotope lower in energy well- escapes less easily
different ZPE =
fractionation during chemical reactions via 2 processes
- equilibrium processes
- kinetic processes
equilibrium processes
rare isotopes do not partition equally between equilibrating species or between different phases of same species at equilibrium
kinetic processes
isotopologues react at different rates in non-reversible reactions
molecular diffusion rates
differ between isotopologues b/c velocity of molecule depends on Ek and inversely on mass
H2 diffuses slightly faster than DH
different diffusion rates =
isotope effects
isotope fractionation
EIE
generally reversible rxn’s or physical processes
governed by ZPE (QMs)
permits isotope exchange
KIE
rate dependent
generally irreversible rxn’s or physical processes
1º - rxn rate determined by rate limiting step
2º -isotopic substitution is remote from from bond being broken
no isotope exchange
e.g. respiration
consequences of mass differences on isotopes
heavier isotope molecules have lower mobility
heavier isotope molecules have higher bonding energy
kinetic energy of a molecule
kT = 1/2 mv^2
molecules have same 1/2mv^2 regardless of isotope content, heavy isotopes have lower v, react slower
determined by temperature
internal energy of gas molecules due to
translational energies
rotational energies
vibrational energies
internal energy of liquid, solid
stretching
vibrational frequency of bonds
12C18O bond strength
higher mass = low vibrational frequency = stronger bond
heavier isotopes have
lower mobility
higher bond energy
heavier isotope species tend to be
concentrated in more dense/strongly bonded phase
isotope fractionation between two phases
decreases w/ increasing T
isotope fractionation factor
α = Ra/Rb
ε =
(α-1)*1000
α_A-B is
ratio of rare/abundant isotope ratio of species in equilibrium
if α = 1, distribution of compounds is equal
deviation from 1 is equil. isotope effect
fractionation factors expressed as 10^3lnα because
close approximation to permit fractionation between materials (ε)
value nearly proportional to inverse of T (1/T) at low T (ºK)
lnα varies as
1/T in low T
1/T^2 in high T
KIE typically
unidirectional A–> B
incomplete, not all of A reacted to B
relatively rapid
KIE examples
Evaporation
diffusion
enzymatic fixation
e.g. CO2 from atmosphere – plant; restricted by stoma
KE
kinetic energy, KE = 0.5mv^2
diffusion ratio
inversely proportional to mass
Da/Db = mb/ma
e.g. 12C16O2 diffuses 1.1% faster than 13C16O2
reduced mass
µ = (mi*M/mi + M)
why reduced mass used
collisions and interactions lower true diffusive rate
reduced mass diffusion ratio
“true ratio”
D1/D2 = sqr. root (µ2/µ1)
Dissociation energy
heavier molecules have higher dissociation energy
therefore bonds harder to break
For reactions that have not reached equilibrium or completion
light isotope is preferentially in the product pool
heavy isotope is in the reactant pool
evaporation under 100% humidity
almost equivalent to evaporation under closed-system conditions
condensation described by
Rayleigh Distillation Rvap = R^o vap * f^(α-1) Rvap = isotope ratio of remaining vapour R^o vap = isotope ratio of initial vapor f = fraction of vapour remaining α = isotopic fractionation factor
Rvap of condensation will be different for Oxygen than H why and how
𝛿18O will be less negative than 𝛿D because of the mass differences
D much greater than H so isotope effects are greater
single-phase, open-system evaporation under equilibrium
𝛿18O of remaining water - increasing exponentially (becoming heavier, more positive) - preferential loss of light
𝛿18O of instantaneous vapour being removed - also increasing exponentially but α lower than 𝛿18O of remaining water ( - at start)
𝛿18O of accumulated vapour being removed - also increasing but much slower/lower sloped
two-phase, closed system evaporation
𝛿18O of water and vapour increase but much less than open system
𝛿18O instantaneous and cumulative vapour identical
isotopic ratios on land
dependent on distance of transport
what kind of process is evaporation
kinetic process
assumptions of evaporation being a kinetic process
rapid
vapour carried away
D
deuterium
fundamental to understanding isotope systematics of hydrologic cycle
knowledge of isotope effects associated with evap and condensation between air masses, reservoirs
condensation is what kind of process
equilibrium process
easier to deal with mathematically, depends on T alone
larger f in Arctic water or equatorial?
Arctic
higher temperature = lower f
higher temperature = lower alpha
isotope effects in open system clouds, assumptions
isotope equilibrium established btw vapour and condensate in cloud
condensate removed from cloud as precip.; no other sources or sinks - cloud is closed; condensate enriched in 18O, D compared to vapour
If cloud undergoes condensate loss under equilibrium conditions and no exchange with environment, change in isotope ratio of remaining vapour
described by Rayleigh Distillation equation for closed system
Rt/Ro = f^(α-1)
f = fraction of cloud vapour remaining
α = Rcondensate/Rvapour
Rayleigh fractionation curve of cloud
𝛿18O vs cloud T and fraction of remaining water
𝛿18O = 0 at top, decreasing down - lighter, more negative
condensate and vapour lines
cloud T changes, dependent on height, alpha dependent on T
larger isotope effects in clouds why
colder T than land where evaporation occurred
latitudinal variation in precipitation
15ºN: 𝛿18O = -2, 𝛿D = -6
25ºN: 𝛿18O=-5, 𝛿D = -30
60ºN: 𝛿18O = -15, 𝛿D = -110
areas of similar rain composition on a map
isoisotopic lines
isotopes in rain controlled by
latitude (# of rain events)
altitude (T)
distance from coast (# of rain events)
GMWL
global meteoric water line 𝛿Dsmow vs 𝛿18Osmow 𝛿D = 8𝛿18O + 10 y axis = 0 down to -500 x axis = -50 right to 0
Hydrogen isotope uses
hydrology
water
mineralogy
geothermometry
Helium isotope uses
mantle, subsurface geochemistry
pathway tracer
Carbon isotope uses
life
biology
partitioning or organic/inorganic compounds, pools
geothermometry
Nitrogen isotope uses
life
trophic levels
biological processes
exogenic carbon cycle
outside of Earths interior
recycled at surface
Major crustal carbon reservoirs
organic carbon (life)
continental crystalline rocks (graphite, diamond)
**sedimentary inorganic C / carbonates (limestones)
Amount of carbon in reservoirs
atmosphere 800 PgC (10^15) ocean ca. 35,000PgC land plant ca. 1000PgC sedimentary 5x10^22gC Corg 15x10^21gC crust 7x10^21
∂13C for C reservoirs
atmos: -8.3‰
ocean: TDC=0‰, DOC=-20‰
land plants:-25‰
sedimentary: 0-1‰
organic: -23‰
crust: -6‰
oxygen isotope uses
paleogeoscience hydrology water mineralogy geothermometry
sulphur isotope uses
global and regional redox state
biological (bacterial) processes
strontium isotope uses
Earth history
∂13C surface ocean
1.8‰
∂13C deep ocean
0.6‰
∂13C fossil fuels
-28‰
why is there different isotope range for primary producers
source C: -8‰ diff. between marine/atmosphere
diff. fractionation processes w/i PP: C4/phtyo./C3
C3
Calvin-Benson
‘normal’ plants
cooler, wetter, cloudier climates
C4 plant
Hatch-Slack
evolved for low CO2 in Cenozoic (65Mya)
bright, dry, warm places
more efficient with water, less efficient with light
examples of C4 plants
maize
sugar cane
desert plants
‘background’ CO2
180ppm
plant minimum
80-100ppm
difference between C3, C4
C3 takes CO2 into mesophyll cell and directly into C-B cycle
C4: CO2– mesophyll– bundle sheath cell– C-B cycle
function of bundle sheath
concentrates CO2
CAM
Crassulacean Acid Metabolism plants
what are CAM plants
have C3 and C4 system that are used separately
what is biggest problem plants have
close stomata when dry conditions to eliminate evaporation - suffocate
C4 plants are advantaged in this way because concentrate CO2 - have some stores
how CAM systems work
Night/rain/cloud: stomata open, build of C pool, no risk of dehydration
day/sunny: stomata closed, feed internally on C pool, no risk of suffocation
where are C3 plants
ubiquitous- all aquatic and ancients
high latitudes, cool climates, forests, woodlands, high latitude grasses
∂13C C3
-23 - -33‰
average -26‰
C4 common plants
tropic/warm grasses, spartina (marsh plant)
when C4 is most favourable
p(CO2) less than 500ppm high p(CO2) C4 is less favourable than C3
∂13C C4
-9 - -16‰
average -13‰
higher plants (angiosperms) -10 - -18‰
10-14‰ more enriched in 13C than C3
∂13C CAM plants
-9 - -33‰
Isotopic range of petroleum
bimodal due to presence of C3/C4 plants, more from C3 (aquatic plants), higher in the -20 - -30‰
C3/C4 plant distribution on earth
C3 in polar regions, tundra, conifer/woodland forests (NA, N Europe), tropical/temperate broad/leaved forests
mixed C3/C4: tropic/temperate desert, semi-desert, tropical woodlands
dominant C4: tropic/temperate grassland
∂13C coal
-25‰
∂13C natural gas
-41‰
∂13C petroleum
-28‰
∂13C anthropogenic CO2
-26‰
making atmosphere lighter, more negative
urban atmosphere ∂13C
-12‰
what happens to isotope ratios during burial and decomp
become heavier
C3 plants: -27‰ –burial - soil org matter -27‰ bacterial decay and respiration(+5‰) - soil CO2 -22‰ – (+10‰) — -12‰
observed discrimination (∂13Catm - ∂13Cplant) vs. (CO2)int/(CO2)ext
no exchange (0,0) – full exchange, open stoma
C4 species plot low on y-axis across x-axis
C3 plot ca. 5-25‰ up y-axis, 0.5-1 on x-axis
why is there observed discrimination
plants that can close stomata - use internal reserves- use light reserves first - (CO2)int decreases
note that in (CO2)int/(CO2)ext only internal is changing in short periods of time, atm is ca. stable
∂13C of photosynthesizers
algae: -10 - -22‰
plankton: -18 - -31‰
kelp: -10 - >-20‰
∂13C of photosynthesizes dependent on
pCO2, T, S, pH
source (atmosphere vs water)
cytoplasm ƒ
temperature effects on ∂13C of photosynth.
higher T = lower fractionation
colder water = more CO2
equatorial plants = lower fractionation
CO2 diffusion rates
land plant ∂13C
-32‰ - -22‰
algae ∂13C
-22‰ - -10‰
∂13C source
∂13C(HCO3) = 0‰ (-0.2‰ in water)
∂13C(CO2) = -8‰ in air
aquatic plants likely use HCO3 and CO2
cytoplasm CO2
in aquatic environment CO2 same phase as cytoplasm CO2 - no isotope effect to convert it
CO2 diffusion
slower in water than air
able to be used more efficiently in aquatic plants (similar to differences in C3 vs C4)
marine vs freshwater plants
marine typically enriched
marine plankton ∂13Corg -28 - -17‰
fresh plankton ∂13C -20‰ - -32‰
why difference in marine vs fresh
HCO3 source
HCO3 ratio
∂13C(HCO3)marine = 0‰ ∂13C(HCO3)fresh = -9‰ - -15‰
HCO3 source
ocean - mostly from atmosphere
lake - mostly from groundwater
∂13C(HCO3)freshwater depends on
residence time
source
degree of equilibration
∂13C(HCO3)freshwater residence time
mixing, exchange with atmospheric CO2
∂13C(HCO3)freshwater source
water body size
circulation
temperature
∂13C(HCO3)freshwater degree of equilibration
equilibration w/ ∂13C(HCO3)groundwater
plankton c-isotope dependence on T
largest α at lowest T = strongest isotope effect = lightest/most negative at high/low latitudes
plankton c-isotope dependence on T, why
enzymatic T-dependence (lower alpha at higher T) CO2 solubility (more soluble at lower T)
why does high [CO2] affect plankton ∂13C
high [CO2] = higher diffusion supply to plankton = large fractionation (like C3 plants)
C isotope dependence on trophic level
plants - herbivores - 1ºconsumer - 2ºconsumer
ratio becoming less negative as go –>
e.g. -25‰ -(-18‰) –(-16‰ ) –(-15/-10‰ )
how to estimate T history of sea water and volume of ice caps
use long-term persisting records - fossil limestone (calcite)
Harold Urey
first to O2 isotopes ratios of carbonates to deduce T of carbonate deposition, now cornerstone of paleooceanography
Basic idea of carbonate paleothermometer
O2 isotope fractionation btw calcite and H2O a fn of T
difference in ∂18O values of calcite/water used to determine T of ocean at time of carbonate formation
obstacles to paleothermometry
able to measure ∂18Ocarb precisely ƒ btw calcite/water well calibrated as a fn of T know if formed in equilibrium degree of change over time ∂18O ocean at time of formation
∂18Ocarbonate formation, equilibrium
disequilibrium if there was rapid precipitation
disequilibrium formation of ∂18O carbonate
biogenic vital effect
example of biogenic vital effect
plankton bloom (growth outside of equilibrium)
∂18Ocalcite change with time
post-burial isotopic exchange w/ pore water
dissolution, recrystallization, etc.
paleothermometry methodology
liberate CO2 from CaCO3 by dissolution w/ H3PO4
number of CO2 isotopologues
12
most common CO2 isotopologues
12C16O16O
13C16O16O
12C18O16O
CO2 liberation from CaCO3 equation
3CaCO3 + 2H3PO4 -> 3CO2 + 3H2O + Ca3(PO4)2
12C16O16O
mass 44
most common isotopologue
98.450 mole%
13C16O16O
mass 45
1.065 mole %
12C18O16O
mass 46
0.405 mole %
C/O isotopes in carbonates referred to
VPDB
O isotopes in water referred to
VSMOW
carbonate EIE
O2 in bicarbonate equilibrates isotopically with O2 in water
bicarbonate used by organisms w/ shells
carbonate EIE equations
Ca2+ + 2HCO3- CaCO3 + H20 + CO2
H2O + CO2 2HCO3-
bicarbonate isotope fractionation factor
α_c-w = Rcalcite/Rwater R_c = 18O/16O in calcite
K of oxygen isotope
K = ([CaC18O3]/[CaC16O3])^1/3 / ([H218O]/[H216O]) K = Rc/Rw = α
Oxygen isotope recorder equation
10000 lnα_c-w = 2.78(T^-2 x 10^6) - 3.39 (T in K)
calcite-aragonite fractionation
grow at same time, same place
∆cal-w vs T = diff. relationship for cal/arag. due to diff. α
use diff. btw relationships to determine T
why use Calcite-Aragonite relationship to determine T
allows you to get around needing to know H2O characteristics
Arag water calc.
arag calcite
calcite-aragonite offset at high T
converge
lower α = less discrimination
∂18Ocarbonate rule of thumb
at constant ∂18Oseawater:
change in ∂18Ocarb of ca. 1% corresponds to ca. 4ºC change
using ∂18Ocarb rule of thumb
use to find T of calcite precipitation only if ratio of water is known
∂18Ocalcite = ∂18Owater - 0.23∆T
using foraminifera
planktonic = surface water T
benthic formas = deep water T = long term/ background/ 1000yr scale/ baseline
benthic foraminifera
18O enriched (cold water)
temperature record from belemnite shell growth rings
can see inter-annual seasonal variability
comparing planktonic to benthic foraminifera
benthic = colder = heavier ∂18O
planktonic = warmer = light/depleted ∂18O
use offset to understand changes in T
carbonate paleotemperature equation variables
∂18Ocarbonate
∂18Owater
Temperature
how T is found for carbonate paleotemp. equation
estimate T from measured ∂18Ocarb assuming ∂18Owater value
validity of T estimate in carbonate paleothermometer depends on
∂18Owater at time of calcite growth
∂18Ocarbonate alteration
carbonate precipitation equilibrium with water
∂18Owater at time of calcite growth
ocean changes due to glacial/interglacial
values buffered by hydrothermal interaction w/ seafloor
shallow epicontinental seas/restricted basins may be very different from deep ocean
∂18Ocarbonate alteration
metamorphism
low T diagenesis
how can we tell if there is ∂18Ocarb alteration
thin sections!
carbonate precipitate in equilibrium with water?
vital effect
some organisms secrete carbonate not in equilibrium
∂18Owater today
30±15‰
ice caps/glaciers = 2.5% of hydrosphere
variation in ∂18Owater
ca. 1.2% between glacial max and interglacial
0. 1‰ / 10m
when ice sheets are more expansive ∂18O
more positive than 0‰
∂18Owater error of 0.2‰
corresponds to T error of 1ºC
glacial ice can be estimated by
sea level changes
sea level, glacial ice change used
make estimates of increase in ocean ∂18Owater during glacial times
problem with using glacial ice change to estimate changes in ∂18Owater
does not consider sea ice (floating, no effect on sea level)
sea ice ratio compared to water
higher 18O/16O ratio
sea level at last glacial extent
-120m
whats happening to ∂18Owater now
adding light water due to glacial melt
foram isotope curves from caribbean cores
show regular changes in sea level associated w/ glacial/interglacial periods
tropical cores less affected by climate change
high latitude snow
light
heavy 18O rich water condenses on mid-lats
atmos. water vapour increasingly depleted in H and O
Interior Antarctica 18O
5% lower than ocean water - meltwater from glacial ice is 18O depleted
interglacial sea level (today)
mean ocean depth = 4000m
∂18O = 0‰ (SMOW)
glacial sea level
∆sea level = 120m
120m/4000m = 0.03 (3% of oceans water frozen)
∆∂18Owater = 1.2‰
why is glacial period sea water ∂18O = 1.2‰
∆∂18O difference between ocean (0‰) and ice (-40‰)
0.0‰ - (-40‰ x 0.03) = 1.2‰
why important to know ∂18O = 1.2‰ during glacial
organisms growing there will have very different isotopic signatures - must correct for the ∆
how to correct for the ∆∂18O from glacial times
using benthic water
∂18O rich carbonate =
colder T
main paleoceanography question
hw much of ∂18O shift is due to ice volume (sea level change) and how much due to T change
LGM
last glacial max
LGM amount of water in ice caps/glaciers
6.5%
∂18O changes with glacial advance on short time scales
ca. 100,000yr entire ocean mass increase ca. 1‰ as light water transferred to ice sheet
marine carbonates change in accordance
what causes changes in carbonate signature during glacial extent/retreat
isotopic shift of water
change in water temperature
Effect of T increase on 18O precipitation
increased T = lower ƒ_carb-wat
organisms precipitate carbonate w/ lower ∂18O values
Effect of T increase on 18O precipitation is compounded by
melting ice caps = reintroduced light water
∂18O of carbonates will also decrease for this
effects of changing T and seawater isotopic composition during glacial cycle
complimentary effects
cause ∂18O value of marine carbonates to change in same direction
how to circumvent the complimentary T/isotope effect (glacial periods)
use tropical organisms (little-no T effect)
use benthic organisms (little-no T effect)
cenozoic glacial-interglacial periods
13 stages over last 500ka
heavier isotope species tend to
concentrate in more dense phase
bond more strongly
isotope fractionation between two phases tends to
decrease with increasing temperature
what happens to cloud as it loses moisture
depleted (more negative)
geographic distribution of isotopes in water
lighter/depleted towards poles
hydrogen shows larger range due to mass differences (0- -270‰)
Deuterium excess
d = ∂D - 8∂18O
global meteoric water line
∂D = 8∂18O + deuterium excess ∂D = 8∂18O + 10 d = 10 for GMWL
18O vs 16O mass difference
2amu
2H vs 1H mass difference
1amu
what is deuterium excess
reflects slower movement of H218O vs HDO during diffusion = enrichment of HDO in less strongly bound phase (gas if water evaporation)
LEL
local evaporation line
slope less than 8 as in GMWL
D-excess increase
in response to enhanced moisture recycle as a result of increased evaporation
D-excess decrease
where water is lost by evaporation
differences in d-excess from
varying T, relative humidity, sea surface wind speed
d-excess, Canada
East coast: ca. 13 - 20
West coast: -4 - +4
meteoric water in lebanon, isotopic composition determined by
3 main air masses:
Eastern Europe- humid, cold
Mediterranean sea - warm, rainy
Syrian desert - warm winds
extra factors in lebanon MWL
mount induce high altitude isotope effect
benefits of altitude effects
distinguish groundwater recharged at high altitudes vs low altitude recharge
Eastern mediterranean d-excess
+22 b/c evap. processes at sea-surface occur due to low-humidity air masses of continental origin
GMWL does not explain mediterranean so well, use
MMWL - mediterranean meteoric water line
LMWL
Lebanese meteoric water line
different slope due to secondary evaporation during rainfall
∂D = 7.135∂18O +15.98
air mass flow in mediterranean sea area
continental air off of Europe converges w/ maritime air (low T, d) off of med. sea – fast evaporation – low T, high d clouds – wind blows over sea to high T, d
∂18O vs altitude
∂18O (-9, -5)
altitude (0, 2500)
decreasing, linear
∆∂18O ca. 2‰ / km
processes that shift values from MWL
O isotopes displaced due to exchange w/ volcanic CO2, limestone
H isotopes due to exchange w/ H2S, silicate hydration, clay dehydration
MWL shifts in minewaters
strongly enriched in D, weakly in 18O
fall above MWL
form mixing lines w/ high 18O, 2H fluid
why MWL shifts in minewaters
WRI of basinal brines
leaching of fluid inclusion in crystalline basement rx
precipitation/exchange w/ hydrous minerals
formation of 1H-rich CH4 or H2
also, latitude effects
isotope ratios in sediments are governed by
meteoric water signal
modulated by exchange, loss in subsurface
enriched by substitution w/ minerals in sediment
additional isotopic impacts on sedimentary fluids
compaction further increases 18O, 2H
connate fluid
formation fluid
MWL shift based on T
shifts less at higher T because of reduced isotope effects
Meteoric waters in oilfield brines
∂18O (0, -20)
degrees latitude north (30, 60)
decreasing ca. linear
AB nearly -20‰
high T geochemistry isotope uses
geothermometry
reconstructing ancient hydrothermal system
detecting crustal assimilation in mantle-derived magma
tracing recycled crust in mantle
exploration, exploitation of geothermal resources
gas isotope geothermometry
kinetics of CO2-H20 very fast
continuous re-equilibration
cannot determine Tmax
slowest rxn rate is CO2-CH4
mineral isotope geothermometry
best have large T coefficient
eg. quartz-magnetite
why quartz-feldspar is not an adequate mineral isotope geothermometer
∂18O T gradient is small
feldspar particularly susceptible to isotopic exchange w/ post-formational fluids
stable isotope thermometry main principle
exchange
fractionation/partioning of heavy/light isotopes between coexisting phases
most important role in isotopic fractionation
vibration motion
fractionation between phases m and n defined as
10^3lnα_m-n = (α10^6 / T) + b (at low T)
mineral - water oxygen isotope exchange
anhydrous mineral - water O2 exchange: b = 3.7
hydrous mineral - water O2 exchange = proportional to number of OH bonds in phase
hydrous mineral - water O2 exchange
b = 3.7 * (1 - fraction of OH bonds / all O-bonds)
KAl3Si3O10(OH)1.8F0.2
b = 3.7 (1- (1.8/10))
O2 isotope in gold exploration
18O depleted rock = most alteration due to pluton intrusion
groundwater percolation = hydrothermal mineral formation/deposition
O isotopes map alteration aureole
geothermal gas fluids contain
CO2, CH4, H2, H2Ovapour
a pair of geothermal gas fluids an
be used as geothermometer
distribution of isotopes between components is a function
of temperature
requirements of geothermometry
isotopic equilibrium approached between species
regular T gradient of isotopic fraction factor α large enough to be easily measurable
mixing w/ same chemical species of different origin excluded
isotopic equilibrium achieve in geothermal reservoir is not altered by sampling
how to avoid altering isotopic equilibrium when sampling
slow rate of isotopic exchange to prevent isotopic re-equilibration btw sampling/analysis
well developed isotopic geothermometers
CO2 + 13CH4 -- 13CO2 + CH4 CH3D + H20 -- HDO + CH4 HD + CH4 -- H2 + CH3D HD + H2O -- H2 + HDO CO2 + H218O -- CO18O + H2O
in systems where gas species concentration is too low for isotopic T determination
SO4 + H218O – SO3*18O + H2O
32SO4 + H234S – 34SO + H232S
why is there seasonal variation in the keeling curve
b/c plant activity lowers it every season
the top is the baseline
∂13CO2 (atmosphere) vs year
decreasing
higher (more 13C) in summer due to higher 12C uptake by plants
lighter isotopes pumped into atmosphere (FF)
seasonal oscillations in atmospheric ∂13C based on latitude
highest oscillation in N, decreases down to South Pole b/c of decrease in plant life - no drawdown effect
atmospheric mixing
yr scale
if atmospheric mixing is quick why is isotopic pattern maintained (variable)
because sources/sinks stay the same
N-S atmospheric gradient
actually small even though most emissions are in NH
NH must be taking up large amounts of emissions
average ∂13C of atmospheric CO2 over ocean
-7.5‰
∂13C CO2 variation w/ time for NH and SH at mid-lats
-0.02‰ /yr
∂13C CO2 continental air
-8 - 9‰
urban influence
variations in ∂13C atmospheric CO2 over ocean
latitude
season
time
∂13C CO2 continental air
large variation for variety of factors
eg uptake/release of biospheric CO2, combustion of ff
1980 vs 2008 CO2
1980: 338ppm, -7.6‰
2008: 380ppm, -8.2‰
Is atmospheric ∂13C change only from FF?
no, if it was we’d expect ∂13C to be -9.85
something else is impacting, e.g. deforestation
∂13C atmospheric vs year
relatively stable at ca. 5 until 1800 where it suddenly and dramatically drops off
the organic cycle can be represented by
CO2 + H2O — CH2O + O2
∂13C global exogenic carbon reservoir
-5‰
surface earth average
∂13C inorganic carbon reservoir
1‰
ƒ = 0.77
∂13C organic carbon reservoir
average Corg = -26‰
ƒ = 0.23
as organic carbon reservoir is depleted in 13C
inorganic carbon reservoir is enriched
m_T ∂_T = m_1*∂_1 + …..
must equal global exogenic C reservoir
if no life (no Corg) then Cinorg would = Cexog.
Corg/Ccarb
has been amazing constant over last 3.5Ga
amount of Corg buried, amount of PP relatively constant through time
major extinctions in earth history
Palaeozoic - Cambrian 550Ma Ordovician 440 Ma Devonian 354 Ma P-T 250Ma Triassic 195Ma K-T 65
great oxidation
2.3 Ga
no increase in Corg burial (none in any of the extinctions either)
sulphur isotopes
32S (95.02%)
33S (0.75%)
34S (4.21%)
36S (0.02%)
S standard
VCDT
Vienna Canyon Diablo Troilite
general terrestrial ∂34S
-50 - +50‰
Ocean ∂34S_SO4
+21‰
∂34S_SO4 in past
Mesozoic: +10‰
Palaeozoic: +30‰
exogenic S partitioned into
inorganic S (gypsum, anhydrite) organic S (anaerobic remineralization of OM)
inorganic S
Ca2+ + SO4- + 2H2O – CaSO4*2H2O
alpha = 1.002 (small)
organic S
remineralized by SRB
OM oxidation
Dissimilatory sulphate reduction
alpha = 1.025 (large
SRB
sulphate reducing bacteria
Organic matter oxidation
2CH2O + SO4- – H2S + 2HCO3-
Dissimilatory sulphate reduction
SO4- + 2H+ + 4H2 —- H2S + 4H2O
Sulphur isotope ratios on Earth
inorganics: igneous, volcanics, petroleum, coal = 0 - +40‰
organics: biogenic pyrite, shales, limestones = -50 - 0‰
inorganic C/S isotopes governed by
EIEs
Keq = alpha_ A-B
organic C/S phases governed by
KIEs
eg. Rayleigh eq
R_A = R_o*ƒ^alpha -1
oxygenated S form
sulphate
anaerobic S form
sulphide (pyrite)
normal marine sediments C:S
8:3
euxinic
anoxic
euxinic sediment C:S
ca. 2:1
non-marine sediment C:S
10:0.5
Earth C/S Cycle
Basic Pool: OandA ∂34S=+19‰, ∂13C=1‰
Inorganic: ∂13C_carbonate=1‰, ∂34S_gypsum=19‰
Organic: ∂34S_shale=-16‰, ∂13C_shale= -26‰
C/S enriched in organic phases
12C, 32S
C/S enriched in inorganic phases
13C, 34S
more burial of carbon implications for oxygen
more oxygen because back reaction does not occur
i.e. back rxn: O2 + CH2O — CO2 + H2O
CH2O buried, O2 remains
more carbon burial implications for sulphur
more Corg burial = more net O2 = less pyrite (aerobic conditions)
C/S ratio through time
increase at 400Ma
peak at 300Ma
periods of high ∂34S
characterized by high rates of pyrite deposition
increased burial Corg
higher ∂13C
higher atmospheric O2
oxidized sulphides to So4
lower ∂34S
largest isotope effects in
KIE
especially enzymatic systems
typical range in C isotopes
atmos. -8‰ bulk plants/kerogen: -20 - 30‰ marine organisms: -10 - -20‰ coal/oil: -30 - -20‰ natural gas: -50 - -20‰ archaea: -120 - -50‰ anaerobic methane: -140 - -30‰
maceral
coal building block inertinite vitrinite exinite sapropoels
∂Dvs∂Corg for macerals
∂D (-130, -70)
∂Corp (-25, -23)
isotopic ranges w/i a single piece of coal
periods of low ∂14S
likely had high rates of gypsum deposition
e.g. Permian
why is there isotopic differences in coal
made up of plant pieces - different masses, diff water w/ diff isotope signatures, diff fractionation w/i plant
hydrocarbon generation
cleave terminal position of long chain = methane = small, light molecule not strongly bonded
isotope effect of cracking kerogen
mostly get 12CH4 - easiest to break off
Petroleum formation
T dependent 20-50ºC = bacterial CH4 50-150ºC = CO2, peak at 100 100-200ºC = thermogenic CH4 150-200ºC = H2S
petroleum windown
80-200ºC
CO2, thermogen. CH4 dominant
methane to atmosphere sources (increasing)
mantle gas, geothermal gas, water column, rice paddies, biogas, terrestrial hydrates, atmosphere, bacterial reservoirs, natural gas, coal/gas reservoirs, marine hydrates
over methane sources
hydrothermal vents
covert methane sources
wetlands
cows
wetland methane emissions
ca 115 Tg CH4/yr
21% o total annual emissions
cow emissions
ca. 55-60 kg CH4/cow/yr
biogeochemical carbon cycling
CO2 – methanogenesis – CH4 – methanotrophy
methanogenesis
carbonate reduction
epsilon_c = 49-95
Methanotrophy
aerobic/anaerobic oxidation CH4 - CO2
most reduced form of C
CH4
Measure isotopic ratio, marine system
may be able to tell system it came from
eg marine: %R:%F = 0:100, slope = 1:1
freshwater: %R:%F = 100:0, slope = 1:0.25
Marine ∂D_CH4 =
∂D_H2O = 18O
methyl fermentation ∂D_CH4
ƒ(∂D_H2O-18O) + (1-ƒ)*[(0.75∂D_methyl) + (0.25∂Hydrogen)]
C-D diagram
∂13C_CH4 vs ∂D_CH4 = methane map, signatures of sources
Y shape - bacterial carbonate reduction, bacterial fermentation, geothermal/hydrothermal
where does Atmospheric plot on C-D diagram
above and right of everything else
unique signature due to sink processes
gas hydrates
10,000Gt C
10-1000X natural gas + coal + oil
meta-stable
will emit huge CH4 if released from permafrost melt
Clathrate gun hypothesis
huge change in water mass signature 4-5‰- how to get that much ‘lightening’ - gas hydrate destabilization?
Types of gas hydrate
Type I/II - biogas
Type H - thermogenic
biogas hydrate ∂13C
∂13C -75 - -60‰
typically -65‰
∂13C thermogenic
-35 - -55‰
typically -45‰
hydrate ice worms ∂13Corg
sediments: -23 - -47‰
aragonite -40 - -45‰
Methane diagenetic carbonate
isotopically very light
formed entirely of CH4 - methane oxidation
carbonate pavement precipitated from CH4
ugly brown, porous, inclusions
carbonate ∂13C
diagenetic -20 - -25‰
methanogenic +10 - +26‰
methanotrophic -70 - -35‰
hydrates and slumping
∆T - methane unfreezes - block of hydrate sediment breaks off - slumping/sliding /debris flow - gas plume released
why do hydrate not usually release gas plumes
SRBs graze down methane - we rely on microbes to protect us
disproving crath ray hypothesis
measure ice samples - see major [CH4] increase in younger dryas ca. 11.4kyr - measure ∂13C_CH4 - see no changes!
what is the significance of no change in ∂14C_CH4 in younger dryas
if the [CH4] increase was due to methane hydrate release would expect a significant ‘lightening’ of the isotopic signature - do not!
what could cause the increased [CH4]
no change in source, just more of it - wetland expansion?
Snowball Earth
Cryogenian period, Neoproterozoic era, ca 650-700Ma- twice?
how to get out of snowball earth
Volcanos! - erupting CO2, no life to take it up - accumulates = warming
how to get in to snowball earth
increased CH4 -increased T - increased weathering and CO2 drawdown - CH4 keeps atmosphere hot for short time period - CH4 comes out of atmos = rapid T drop
fate of CO2 mostly dictated by
weathering - ultimate CO2 sink
more weathering = more CO2 sink
∂13C vs age, Meso, Neoproterozoic
leading in to Neoprot. see more highs/lows of ∂13C - transitioning in and out of snowball phases?
∂13C_carb and burial
higher ∂13C_carb = more organic burial
Alkenones
C37 ketones. di/tri-unsaturated long-chain alkenones
using alkenones for paleothermometry
coccolithophores change amount of alkenone in membrane and therefore fluidity of membrane based on T
Where do alkenones come from
uniquely derived by haptophytes (eg. coccolithophores)
important coccolithophore
Emiliania huxleyi
U^k_37 index
[C_37:2]/([C_37:3] + [C_37:3])
degree of unsaturation - fn of SST, specific to E Huxleyi growth
why use alkenone-derive paleo barometry
can tell up to 30Ma, ice cores only tell ca 1Ma
