EOS 335 Part II Flashcards
(383 cards)
1
Q
elements with at least 1 stable isotope
A
80
2
Q
known stable isotopes
A
250
3
Q
element with most stable isotopes
A
Tin, 10, 112Sn - 124Sn
4
Q
elements with 8 stable isotopes
A
none
5
Q
elements with 9 stable isotopes
A
only xenon
6
Q
mononuclidic elements
A
27 (single isotope)
7
Q
elements with at least 1 stable isotope
A
H - Pb (1-82) except technetium and promethium
8
Q
elements without stable isotopes
A
> 82
9
Q
stables isotopes are what state
A
ground state of a nuclei
10
Q
isotopic elements of geochemical/biological interest have
A
2+ stable isotopes
lightest generally in greater abundance
(C,H,O,N,S)
11
Q
isotope properties
A
same protons and electrons = same chemical behaviour
physiochemical differences
12
Q
isotope properties, same chemical behaviour
A
enter same chemical reactions
form same bonds
rare isotope can trace abundant isotope
13
Q
lighter stable isotope
A
generally more abundant
not Li, B
14
Q
Important stable isotopes
A
H, D, C, N, O, S, Cl
15
Q
isotope properties, physiochemical differences
A
lead to differences in distribution between phases boiling pt freezing pt density vapor pressure, etc.
16
Q
Characteristics of elements for isotope effects
A
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
17
Q
low atomic mass, isotope effects
A
H, He, C, N, O, S, Cl
exception - Fe isotopes fractionated by bacteria
18
Q
large mass difference, isotope effects
A
∆m D-H = ca. 100%
∆m 13C-12C = 8.3%
∆m 18O-16O = 12.5%
19
Q
why measure stable isotopes as ratios
A
utility- compare identical species/phases
measurement - measuring ratios increases precision
20
Q
δ
A
‰
unitless
differences between sample and standard readings
not absolute isotope abundance
21
Q
natural abundance standard
A
defined as δ = 0
22
Q
Stable isotope ratio notations
A
δ^n X sample (‰) = [Rsample - Rstandard/ Standard] x1000
23
Q
R
A
absolute abundance ratio
atom% ^n X / atom% ^m X
e.g. atom% 15N /atom % 14N
24
Q
light/heavy wording
A
lots of heavy isotope = enriched, heavy less of heavy isotope = depleted, light e.g. more 13C = heavy, enriched, + less 13C = light, depleted, - 0 = standard
25
hydrogen stable isotope standards
SMOW, VSMOW
| standard mean ocean water
26
oxygen stable isotope standards
SMOW VSMOW
| standard mean ocean water
27
carbon stable isotope standards
PDB VPDB
Pee Dee Belemnite
13C/12C = 0.0112372
18O/16O = 0.0020671
28
nitrogen stable isotope standards
atm N2
| atmospheric nitrogen
29
sulphur stable isotope ratio
CDT
| Canyon diablo triolite
30
δ13C atmospheric CO2
-8.2‰
31
δ13C plants, kerogen, coal
-8 - -55‰
32
𝛿 13C oil
-22 - -50‰
33
𝛿13C natural gas
-25 - -100‰
34
range in hydrogen isotopic variation
-700 - +200‰
35
range in carbon isotopic variation
-140 - +40‰
36
range in nitrogen isotopic variation
-60 - +50‰
37
range in oxygen isotopic variation
-30 - +30‰
38
range in sulphur isotopic variation
-50 - +40‰
39
Hydrogen isotope mass differences
99.8%
40
Carbon isotope mass differences
8.36%
41
nitrogen isotope mass differences
7.12%
42
isotopes are subjected to
Isotope effects
43
higher mass differences =
larger isotope effects
| more strongly fractionated
44
Isotope effects
departure in isotope ratio from global average abundance due to physiochemical mechanisms
45
types of isotope effects
EIE - equilibrium isotope effects
| KIE - kinetic isotope effects
46
Isotope fractionation
expression of isotope effects
47
Oxygen mass differences
∆m 16O/18O = 12.%
48
sulphur isotope mass differences
∆m 32S/34S = 6.24%
49
causes for isotope effects
chemical, physical properties of isotope
physiochemical properties of isotopes
isotopologues
50
chemical, physical properties of isotopes
arise from differences in atomic mass
reaction rates
diffusion rates
equilibrium constants
51
physiochemical properties of isotopes
result of quantum mechanical effects
energy of molecule restricted to discrete E levels
e.g. heat capacity, density, vapour pressure
52
isotopologues of isotopes
same molecules with different masses
different vibrational energies
e.g. H2, DH, D2
53
different isotopologues
different masses = different vibrational energies = different zero-point energies
54
ZPE
zero point energy
| energy difference between minimum in potential energy curve and ground state energy (Eo)
55
Eo =
1/2 hv
h = plancks constant
v = vibrational frequency
56
ZPE, heavy isotopologue
lower ZPE than lighter isotopologue because v varies inversely with mass
57
isotopologue bonds
weaker in light isotopologues, easier to break, due to higher ZPE
58
difference in chemical properties of isotopologues can be evaluated in terms of
vibrational frequencies
| transition states, in case of kinetic effects
59
energy well, isotopologues
heavier isotope lower in energy well- escapes less easily
60
different ZPE =
fractionation during chemical reactions via 2 processes
- equilibrium processes
- kinetic processes
61
equilibrium processes
rare isotopes do not partition equally between equilibrating species or between different phases of same species at equilibrium
62
kinetic processes
isotopologues react at different rates in non-reversible reactions
63
molecular diffusion rates
differ between isotopologues b/c velocity of molecule depends on Ek and inversely on mass
H2 diffuses slightly faster than DH
64
different diffusion rates =
isotope effects
| isotope fractionation
65
EIE
generally reversible rxn's or physical processes
governed by ZPE (QMs)
permits isotope exchange
66
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
67
consequences of mass differences on isotopes
heavier isotope molecules have lower mobility
| heavier isotope molecules have higher bonding energy
68
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
69
internal energy of gas molecules due to
translational energies
rotational energies
vibrational energies
70
internal energy of liquid, solid
stretching
| vibrational frequency of bonds
71
12C18O bond strength
higher mass = low vibrational frequency = stronger bond
72
heavier isotopes have
lower mobility
| higher bond energy
73
heavier isotope species tend to be
concentrated in more dense/strongly bonded phase
74
isotope fractionation between two phases
decreases w/ increasing T
75
isotope fractionation factor
α = Ra/Rb
76
ε =
(α-1)*1000
77
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
78
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)
79
lnα varies as
1/T in low T
| 1/T^2 in high T
80
KIE typically
unidirectional A--> B
incomplete, not all of A reacted to B
relatively rapid
81
KIE examples
Evaporation
diffusion
enzymatic fixation
e.g. CO2 from atmosphere -- plant; restricted by stoma
82
KE
kinetic energy, KE = 0.5mv^2
83
diffusion ratio
inversely proportional to mass
Da/Db = mb/ma
e.g. 12C16O2 diffuses 1.1% faster than 13C16O2
84
reduced mass
µ = (mi*M/mi + M)
85
why reduced mass used
collisions and interactions lower true diffusive rate
86
reduced mass diffusion ratio
"true ratio"
| D1/D2 = sqr. root (µ2/µ1)
87
Dissociation energy
heavier molecules have higher dissociation energy
| therefore bonds harder to break
88
For reactions that have not reached equilibrium or completion
light isotope is preferentially in the product pool
| heavy isotope is in the reactant pool
89
evaporation under 100% humidity
almost equivalent to evaporation under closed-system conditions
90
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
```
91
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
92
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
93
two-phase, closed system evaporation
𝛿18O of water and vapour increase but much less than open system
𝛿18O instantaneous and cumulative vapour identical
94
isotopic ratios on land
dependent on distance of transport
95
what kind of process is evaporation
kinetic process
96
assumptions of evaporation being a kinetic process
rapid
| vapour carried away
97
D
deuterium
98
fundamental to understanding isotope systematics of hydrologic cycle
knowledge of isotope effects associated with evap and condensation between air masses, reservoirs
99
condensation is what kind of process
equilibrium process
| easier to deal with mathematically, depends on T alone
100
larger f in Arctic water or equatorial?
Arctic
higher temperature = lower f
higher temperature = lower alpha
101
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
102
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
103
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
104
larger isotope effects in clouds why
colder T than land where evaporation occurred
105
latitudinal variation in precipitation
15ºN: 𝛿18O = -2, 𝛿D = -6
25ºN: 𝛿18O=-5, 𝛿D = -30
60ºN: 𝛿18O = -15, 𝛿D = -110
106
areas of similar rain composition on a map
isoisotopic lines
107
isotopes in rain controlled by
latitude (# of rain events)
altitude (T)
distance from coast (# of rain events)
108
GMWL
```
global meteoric water line
𝛿Dsmow vs 𝛿18Osmow
𝛿D = 8𝛿18O + 10
y axis = 0 down to -500
x axis = -50 right to 0
```
109
Hydrogen isotope uses
hydrology
water
mineralogy
geothermometry
110
Helium isotope uses
mantle, subsurface geochemistry
| pathway tracer
111
Carbon isotope uses
life
biology
partitioning or organic/inorganic compounds, pools
geothermometry
112
Nitrogen isotope uses
life
trophic levels
biological processes
113
exogenic carbon cycle
outside of Earths interior
| recycled at surface
114
Major crustal carbon reservoirs
organic carbon (life)
continental crystalline rocks (graphite, diamond)
**sedimentary inorganic C / carbonates (limestones)
115
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
```
116
∂13C for C reservoirs
atmos: -8.3‰
ocean: TDC=0‰, DOC=-20‰
land plants:-25‰
sedimentary: 0-1‰
organic: -23‰
crust: -6‰
117
oxygen isotope uses
```
paleogeoscience
hydrology
water
mineralogy
geothermometry
```
118
sulphur isotope uses
global and regional redox state
| biological (bacterial) processes
119
strontium isotope uses
Earth history
120
∂13C surface ocean
1.8‰
121
∂13C deep ocean
0.6‰
122
∂13C fossil fuels
-28‰
123
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
124
C3
Calvin-Benson
'normal' plants
cooler, wetter, cloudier climates
125
C4 plant
Hatch-Slack
evolved for low CO2 in Cenozoic (65Mya)
bright, dry, warm places
more efficient with water, less efficient with light
126
examples of C4 plants
maize
sugar cane
desert plants
127
'background' CO2
180ppm
128
plant minimum
80-100ppm
129
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
130
function of bundle sheath
concentrates CO2
131
CAM
Crassulacean Acid Metabolism plants
132
what are CAM plants
have C3 and C4 system that are used separately
133
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
134
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
135
where are C3 plants
ubiquitous- all aquatic and ancients
| high latitudes, cool climates, forests, woodlands, high latitude grasses
136
∂13C C3
-23 - -33‰
| average -26‰
137
C4 common plants
tropic/warm grasses, spartina (marsh plant)
138
when C4 is most favourable
```
p(CO2) less than 500ppm
high p(CO2) C4 is less favourable than C3
```
139
∂13C C4
-9 - -16‰
average -13‰
higher plants (angiosperms) -10 - -18‰
10-14‰ more enriched in 13C than C3
140
∂13C CAM plants
-9 - -33‰
141
Isotopic range of petroleum
bimodal due to presence of C3/C4 plants, more from C3 (aquatic plants), higher in the -20 - -30‰
142
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
143
∂13C coal
-25‰
144
∂13C natural gas
-41‰
145
∂13C petroleum
-28‰
146
∂13C anthropogenic CO2
-26‰
| making atmosphere lighter, more negative
147
urban atmosphere ∂13C
-12‰
148
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‰
149
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
150
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
151
∂13C of photosynthesizers
algae: -10 - -22‰
plankton: -18 - -31‰
kelp: -10 - >-20‰
152
∂13C of photosynthesizes dependent on
pCO2, T, S, pH
source (atmosphere vs water)
cytoplasm ƒ
153
temperature effects on ∂13C of photosynth.
higher T = lower fractionation
colder water = more CO2
equatorial plants = lower fractionation
CO2 diffusion rates
154
land plant ∂13C
-32‰ - -22‰
155
algae ∂13C
-22‰ - -10‰
156
∂13C source
∂13C(HCO3) = 0‰ (-0.2‰ in water)
∂13C(CO2) = -8‰ in air
aquatic plants likely use HCO3 and CO2
157
cytoplasm CO2
in aquatic environment CO2 same phase as cytoplasm CO2 - no isotope effect to convert it
158
CO2 diffusion
slower in water than air
| able to be used more efficiently in aquatic plants (similar to differences in C3 vs C4)
159
marine vs freshwater plants
marine typically enriched
marine plankton ∂13Corg -28 - -17‰
fresh plankton ∂13C -20‰ - -32‰
160
why difference in marine vs fresh
HCO3 source
161
HCO3 ratio
```
∂13C(HCO3)marine = 0‰
∂13C(HCO3)fresh = -9‰ - -15‰
```
162
HCO3 source
ocean - mostly from atmosphere
| lake - mostly from groundwater
163
∂13C(HCO3)freshwater depends on
residence time
source
degree of equilibration
164
∂13C(HCO3)freshwater residence time
mixing, exchange with atmospheric CO2
165
∂13C(HCO3)freshwater source
water body size
circulation
temperature
166
∂13C(HCO3)freshwater degree of equilibration
equilibration w/ ∂13C(HCO3)groundwater
167
plankton c-isotope dependence on T
largest α at lowest T = strongest isotope effect = lightest/most negative at high/low latitudes
168
plankton c-isotope dependence on T, why
```
enzymatic T-dependence (lower alpha at higher T)
CO2 solubility (more soluble at lower T)
```
169
why does high [CO2] affect plankton ∂13C
high [CO2] = higher diffusion supply to plankton = large fractionation (like C3 plants)
170
C isotope dependence on trophic level
plants - herbivores - 1ºconsumer - 2ºconsumer
ratio becoming less negative as go -->
e.g. -25‰ -(-18‰) --(-16‰ ) --(-15/-10‰ )
171
how to estimate T history of sea water and volume of ice caps
use long-term persisting records - fossil limestone (calcite)
172
Harold Urey
first to O2 isotopes ratios of carbonates to deduce T of carbonate deposition, now cornerstone of paleooceanography
173
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
174
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
```
175
∂18Ocarbonate formation, equilibrium
disequilibrium if there was rapid precipitation
176
disequilibrium formation of ∂18O carbonate
biogenic vital effect
177
example of biogenic vital effect
plankton bloom (growth outside of equilibrium)
178
∂18Ocalcite change with time
post-burial isotopic exchange w/ pore water
| dissolution, recrystallization, etc.
179
paleothermometry methodology
liberate CO2 from CaCO3 by dissolution w/ H3PO4
180
number of CO2 isotopologues
12
181
most common CO2 isotopologues
12C16O16O
13C16O16O
12C18O16O
182
CO2 liberation from CaCO3 equation
3CaCO3 + 2H3PO4 -> 3CO2 + 3H2O + Ca3(PO4)2
183
12C16O16O
mass 44
most common isotopologue
98.450 mole%
184
13C16O16O
mass 45
| 1.065 mole %
185
12C18O16O
mass 46
| 0.405 mole %
186
C/O isotopes in carbonates referred to
VPDB
187
O isotopes in water referred to
VSMOW
188
carbonate EIE
O2 in bicarbonate equilibrates isotopically with O2 in water
bicarbonate used by organisms w/ shells
189
carbonate EIE equations
Ca2+ + 2HCO3- CaCO3 + H20 + CO2
| H2O + CO2 2HCO3-
190
bicarbonate isotope fractionation factor
```
α_c-w = Rcalcite/Rwater
R_c = 18O/16O in calcite
```
191
K of oxygen isotope
```
K = ([CaC18O3]/[CaC16O3])^1/3 / ([H218O]/[H216O])
K = Rc/Rw = α
```
192
Oxygen isotope recorder equation
10000 lnα_c-w = 2.78(T^-2 x 10^6) - 3.39 (T in K)
193
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
194
why use Calcite-Aragonite relationship to determine T
allows you to get around needing to know H2O characteristics
Arag water calc.
arag calcite
195
calcite-aragonite offset at high T
converge
| lower α = less discrimination
196
∂18Ocarbonate rule of thumb
at constant ∂18Oseawater:
| change in ∂18Ocarb of ca. 1% corresponds to ca. 4ºC change
197
using ∂18Ocarb rule of thumb
use to find T of calcite precipitation only if ratio of water is known
∂18Ocalcite = ∂18Owater - 0.23∆T
198
using foraminifera
planktonic = surface water T
| benthic formas = deep water T = long term/ background/ 1000yr scale/ baseline
199
benthic foraminifera
18O enriched (cold water)
200
temperature record from belemnite shell growth rings
can see inter-annual seasonal variability
201
comparing planktonic to benthic foraminifera
benthic = colder = heavier ∂18O
planktonic = warmer = light/depleted ∂18O
use offset to understand changes in T
202
carbonate paleotemperature equation variables
∂18Ocarbonate
∂18Owater
Temperature
203
how T is found for carbonate paleotemp. equation
estimate T from measured ∂18Ocarb assuming ∂18Owater value
204
validity of T estimate in carbonate paleothermometer depends on
∂18Owater at time of calcite growth
∂18Ocarbonate alteration
carbonate precipitation equilibrium with water
205
∂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
206
∂18Ocarbonate alteration
metamorphism
| low T diagenesis
207
how can we tell if there is ∂18Ocarb alteration
thin sections!
208
carbonate precipitate in equilibrium with water?
vital effect
| some organisms secrete carbonate not in equilibrium
209
∂18Owater today
30±15‰
| ice caps/glaciers = 2.5% of hydrosphere
210
variation in ∂18Owater
ca. 1.2% between glacial max and interglacial
| 0. 1‰ / 10m
211
when ice sheets are more expansive ∂18O
more positive than 0‰
212
∂18Owater error of 0.2‰
corresponds to T error of 1ºC
213
glacial ice can be estimated by
sea level changes
214
sea level, glacial ice change used
make estimates of increase in ocean ∂18Owater during glacial times
215
problem with using glacial ice change to estimate changes in ∂18Owater
does not consider sea ice (floating, no effect on sea level)
216
sea ice ratio compared to water
higher 18O/16O ratio
217
sea level at last glacial extent
-120m
218
whats happening to ∂18Owater now
adding light water due to glacial melt
219
foram isotope curves from caribbean cores
show regular changes in sea level associated w/ glacial/interglacial periods
tropical cores less affected by climate change
220
high latitude snow
light
heavy 18O rich water condenses on mid-lats
atmos. water vapour increasingly depleted in H and O
221
Interior Antarctica 18O
5% lower than ocean water - meltwater from glacial ice is 18O depleted
222
interglacial sea level (today)
mean ocean depth = 4000m
| ∂18O = 0‰ (SMOW)
223
glacial sea level
∆sea level = 120m
120m/4000m = 0.03 (3% of oceans water frozen)
∆∂18Owater = 1.2‰
224
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‰
225
why important to know ∂18O = 1.2‰ during glacial
organisms growing there will have very different isotopic signatures - must correct for the ∆
226
how to correct for the ∆∂18O from glacial times
using benthic water
227
∂18O rich carbonate =
colder T
228
main paleoceanography question
hw much of ∂18O shift is due to ice volume (sea level change) and how much due to T change
229
LGM
last glacial max
230
LGM amount of water in ice caps/glaciers
6.5%
231
∂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
232
what causes changes in carbonate signature during glacial extent/retreat
isotopic shift of water
| change in water temperature
233
Effect of T increase on 18O precipitation
increased T = lower ƒ_carb-wat
| organisms precipitate carbonate w/ lower ∂18O values
234
Effect of T increase on 18O precipitation is compounded by
melting ice caps = reintroduced light water
| ∂18O of carbonates will also decrease for this
235
effects of changing T and seawater isotopic composition during glacial cycle
complimentary effects
| cause ∂18O value of marine carbonates to change in same direction
236
how to circumvent the complimentary T/isotope effect (glacial periods)
use tropical organisms (little-no T effect)
| use benthic organisms (little-no T effect)
237
cenozoic glacial-interglacial periods
13 stages over last 500ka
238
heavier isotope species tend to
concentrate in more dense phase
| bond more strongly
239
isotope fractionation between two phases tends to
decrease with increasing temperature
240
what happens to cloud as it loses moisture
depleted (more negative)
241
geographic distribution of isotopes in water
lighter/depleted towards poles
| hydrogen shows larger range due to mass differences (0- -270‰)
242
Deuterium excess
d = ∂D - 8∂18O
243
global meteoric water line
```
∂D = 8∂18O + deuterium excess
∂D = 8∂18O + 10
d = 10 for GMWL
```
244
18O vs 16O mass difference
2amu
245
2H vs 1H mass difference
1amu
246
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)
247
LEL
local evaporation line
| slope less than 8 as in GMWL
248
D-excess increase
in response to enhanced moisture recycle as a result of increased evaporation
249
D-excess decrease
where water is lost by evaporation
250
differences in d-excess from
varying T, relative humidity, sea surface wind speed
251
d-excess, Canada
East coast: ca. 13 - 20
| West coast: -4 - +4
252
meteoric water in lebanon, isotopic composition determined by
3 main air masses:
Eastern Europe- humid, cold
Mediterranean sea - warm, rainy
Syrian desert - warm winds
253
extra factors in lebanon MWL
mount induce high altitude isotope effect
254
benefits of altitude effects
distinguish groundwater recharged at high altitudes vs low altitude recharge
255
Eastern mediterranean d-excess
+22 b/c evap. processes at sea-surface occur due to low-humidity air masses of continental origin
256
GMWL does not explain mediterranean so well, use
MMWL - mediterranean meteoric water line
257
LMWL
Lebanese meteoric water line
different slope due to secondary evaporation during rainfall
∂D = 7.135∂18O +15.98
258
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
259
∂18O vs altitude
∂18O (-9, -5)
altitude (0, 2500)
decreasing, linear
∆∂18O ca. 2‰ / km
260
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
261
MWL shifts in minewaters
strongly enriched in D, weakly in 18O
fall above MWL
form mixing lines w/ high 18O, 2H fluid
262
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
263
isotope ratios in sediments are governed by
meteoric water signal
modulated by exchange, loss in subsurface
enriched by substitution w/ minerals in sediment
264
additional isotopic impacts on sedimentary fluids
compaction further increases 18O, 2H
265
connate fluid
formation fluid
266
MWL shift based on T
shifts less at higher T because of reduced isotope effects
267
Meteoric waters in oilfield brines
∂18O (0, -20)
degrees latitude north (30, 60)
decreasing ca. linear
AB nearly -20‰
268
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
269
gas isotope geothermometry
kinetics of CO2-H20 very fast
continuous re-equilibration
cannot determine Tmax
slowest rxn rate is CO2-CH4
270
mineral isotope geothermometry
best have large T coefficient
| eg. quartz-magnetite
271
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
272
stable isotope thermometry main principle
exchange
| fractionation/partioning of heavy/light isotopes between coexisting phases
273
most important role in isotopic fractionation
vibration motion
274
fractionation between phases m and n defined as
10^3lnα_m-n = (α10^6 / T) + b (at low T)
275
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
276
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))
277
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
278
geothermal gas fluids contain
CO2, CH4, H2, H2Ovapour
279
a pair of geothermal gas fluids an
be used as geothermometer
280
distribution of isotopes between components is a function
of temperature
281
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
282
how to avoid altering isotopic equilibrium when sampling
slow rate of isotopic exchange to prevent isotopic re-equilibration btw sampling/analysis
283
well developed isotopic geothermometers
```
CO2 + 13CH4 -- 13CO2 + CH4
CH3D + H20 -- HDO + CH4
HD + CH4 -- H2 + CH3D
HD + H2O -- H2 + HDO
CO2 + H218O -- CO18O + H2O
```
284
in systems where gas species concentration is too low for isotopic T determination
SO4 + H218O -- SO3*18O + H2O
| 32SO4 + H2*34S -- 34SO + H2*32S
285
why is there seasonal variation in the keeling curve
b/c plant activity lowers it every season
| the top is the baseline
286
∂13CO2 (atmosphere) vs year
decreasing
higher (more 13C) in summer due to higher 12C uptake by plants
lighter isotopes pumped into atmosphere (FF)
287
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
288
atmospheric mixing
yr scale
289
if atmospheric mixing is quick why is isotopic pattern maintained (variable)
because sources/sinks stay the same
290
N-S atmospheric gradient
actually small even though most emissions are in NH
| NH must be taking up large amounts of emissions
291
average ∂13C of atmospheric CO2 over ocean
-7.5‰
292
∂13C CO2 variation w/ time for NH and SH at mid-lats
-0.02‰ /yr
293
∂13C CO2 continental air
-8 - 9‰
| urban influence
294
variations in ∂13C atmospheric CO2 over ocean
latitude
season
time
295
∂13C CO2 continental air
large variation for variety of factors
| eg uptake/release of biospheric CO2, combustion of ff
296
1980 vs 2008 CO2
1980: 338ppm, -7.6‰
2008: 380ppm, -8.2‰
297
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
298
∂13C atmospheric vs year
relatively stable at ca. 5 until 1800 where it suddenly and dramatically drops off
299
the organic cycle can be represented by
CO2 + H2O --- CH2O + O2
300
∂13C global exogenic carbon reservoir
-5‰
| surface earth average
301
∂13C inorganic carbon reservoir
1‰
| ƒ = 0.77
302
∂13C organic carbon reservoir
average Corg = -26‰
| ƒ = 0.23
303
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.
304
Corg/Ccarb
has been amazing constant over last 3.5Ga
| amount of Corg buried, amount of PP relatively constant through time
305
major extinctions in earth history
```
Palaeozoic - Cambrian 550Ma
Ordovician 440 Ma
Devonian 354 Ma
P-T 250Ma
Triassic 195Ma
K-T 65
```
306
great oxidation
2.3 Ga
| no increase in Corg burial (none in any of the extinctions either)
307
sulphur isotopes
32S (95.02%)
33S (0.75%)
34S (4.21%)
36S (0.02%)
308
S standard
VCDT
| Vienna Canyon Diablo Troilite
309
general terrestrial ∂34S
-50 - +50‰
310
Ocean ∂34S_SO4
+21‰
311
∂34S_SO4 in past
Mesozoic: +10‰
Palaeozoic: +30‰
312
exogenic S partitioned into
```
inorganic S (gypsum, anhydrite)
organic S (anaerobic remineralization of OM)
```
313
inorganic S
Ca2+ + SO4- + 2H2O -- CaSO4*2H2O
| alpha = 1.002 (small)
314
organic S
remineralized by SRB
OM oxidation
Dissimilatory sulphate reduction
alpha = 1.025 (large
315
SRB
sulphate reducing bacteria
316
Organic matter oxidation
2CH2O + SO4- -- H2S + 2HCO3-
317
Dissimilatory sulphate reduction
SO4- + 2H+ + 4H2 ---- H2S + 4H2O
318
Sulphur isotope ratios on Earth
inorganics: igneous, volcanics, petroleum, coal = 0 - +40‰
organics: biogenic pyrite, shales, limestones = -50 - 0‰
319
inorganic C/S isotopes governed by
EIEs
| Keq = alpha_ A-B
320
organic C/S phases governed by
KIEs
eg. Rayleigh eq
R_A = R_o*ƒ^alpha -1
321
oxygenated S form
sulphate
322
anaerobic S form
sulphide (pyrite)
323
normal marine sediments C:S
8:3
324
euxinic
anoxic
325
euxinic sediment C:S
ca. 2:1
326
non-marine sediment C:S
10:0.5
327
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‰
328
C/S enriched in organic phases
12C, 32S
329
C/S enriched in inorganic phases
13C, 34S
330
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
331
more carbon burial implications for sulphur
more Corg burial = more net O2 = less pyrite (aerobic conditions)
332
C/S ratio through time
increase at 400Ma
| peak at 300Ma
333
periods of high ∂34S
characterized by high rates of pyrite deposition
334
increased burial Corg
higher ∂13C
higher atmospheric O2
oxidized sulphides to So4
lower ∂34S
335
largest isotope effects in
KIE
| especially enzymatic systems
336
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‰
```
337
maceral
```
coal building block
inertinite
vitrinite
exinite
sapropoels
```
338
∂Dvs∂Corg for macerals
∂D (-130, -70)
∂Corp (-25, -23)
isotopic ranges w/i a single piece of coal
339
periods of low ∂14S
likely had high rates of gypsum deposition
| e.g. Permian
340
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
341
hydrocarbon generation
cleave terminal position of long chain = methane = small, light molecule not strongly bonded
342
isotope effect of cracking kerogen
mostly get 12CH4 - easiest to break off
343
Petroleum formation
```
T dependent
20-50ºC = bacterial CH4
50-150ºC = CO2, peak at 100
100-200ºC = thermogenic CH4
150-200ºC = H2S
```
344
petroleum windown
80-200ºC
| CO2, thermogen. CH4 dominant
345
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
346
over methane sources
hydrothermal vents
347
covert methane sources
wetlands
| cows
348
wetland methane emissions
ca 115 Tg CH4/yr
| 21% o total annual emissions
349
cow emissions
ca. 55-60 kg CH4/cow/yr
350
biogeochemical carbon cycling
CO2 -- methanogenesis -- CH4 -- methanotrophy
351
methanogenesis
carbonate reduction
| epsilon_c = 49-95
352
Methanotrophy
aerobic/anaerobic oxidation CH4 - CO2
353
most reduced form of C
CH4
354
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
355
Marine ∂D_CH4 =
∂D_H2O = 18O
356
methyl fermentation ∂D_CH4
ƒ(∂D_H2O-18O) + (1-ƒ)*[(0.75∂D_methyl) + (0.25∂Hydrogen)]
357
C-D diagram
∂13C_CH4 vs ∂D_CH4 = methane map, signatures of sources
| Y shape - bacterial carbonate reduction, bacterial fermentation, geothermal/hydrothermal
358
where does Atmospheric plot on C-D diagram
above and right of everything else
| unique signature due to sink processes
359
gas hydrates
10,000Gt C
10-1000X natural gas + coal + oil
meta-stable
will emit huge CH4 if released from permafrost melt
360
Clathrate gun hypothesis
huge change in water mass signature 4-5‰- how to get that much 'lightening' - gas hydrate destabilization?
361
Types of gas hydrate
Type I/II - biogas
| Type H - thermogenic
362
biogas hydrate ∂13C
∂13C -75 - -60‰
| typically -65‰
363
∂13C thermogenic
-35 - -55‰
| typically -45‰
364
hydrate ice worms ∂13Corg
sediments: -23 - -47‰
| aragonite -40 - -45‰
365
Methane diagenetic carbonate
isotopically very light
formed entirely of CH4 - methane oxidation
carbonate pavement precipitated from CH4
ugly brown, porous, inclusions
366
carbonate ∂13C
diagenetic -20 - -25‰
methanogenic +10 - +26‰
methanotrophic -70 - -35‰
367
hydrates and slumping
∆T - methane unfreezes - block of hydrate sediment breaks off - slumping/sliding /debris flow - gas plume released
368
why do hydrate not usually release gas plumes
SRBs graze down methane - we rely on microbes to protect us
369
disproving crath ray hypothesis
measure ice samples - see major [CH4] increase in younger dryas ca. 11.4kyr - measure ∂13C_CH4 - see no changes!
370
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!
371
what could cause the increased [CH4]
no change in source, just more of it - wetland expansion?
372
Snowball Earth
Cryogenian period, Neoproterozoic era, ca 650-700Ma- twice?
373
how to get out of snowball earth
Volcanos! - erupting CO2, no life to take it up - accumulates = warming
374
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
375
fate of CO2 mostly dictated by
weathering - ultimate CO2 sink
| more weathering = more CO2 sink
376
∂13C vs age, Meso, Neoproterozoic
leading in to Neoprot. see more highs/lows of ∂13C - transitioning in and out of snowball phases?
377
∂13C_carb and burial
higher ∂13C_carb = more organic burial
378
Alkenones
C37 ketones. di/tri-unsaturated long-chain alkenones
379
using alkenones for paleothermometry
coccolithophores change amount of alkenone in membrane and therefore fluidity of membrane based on T
380
Where do alkenones come from
uniquely derived by haptophytes (eg. coccolithophores)
381
important coccolithophore
Emiliania huxleyi
382
U^k_37 index
[C_37:2]/([C_37:3] + [C_37:3])
| degree of unsaturation - fn of SST, specific to E Huxleyi growth
383
why use alkenone-derive paleo barometry
can tell up to 30Ma, ice cores only tell ca 1Ma
