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2016/2017 > EOS 335 Part II > Flashcards

EOS 335 Part II Flashcards

1
Q

elements with at least 1 stable isotope

A

80

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2
Q

known stable isotopes

A

250

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3
Q

element with most stable isotopes

A

Tin, 10, 112Sn - 124Sn

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4
Q

elements with 8 stable isotopes

A

none

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5
Q

elements with 9 stable isotopes

A

only xenon

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6
Q

mononuclidic elements

A

27 (single isotope)

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7
Q

elements with at least 1 stable isotope

A

H - Pb (1-82) except technetium and promethium

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8
Q

elements without stable isotopes

A

> 82

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9
Q

stables isotopes are what state

A

ground state of a nuclei

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10
Q

isotopic elements of geochemical/biological interest have

A

2+ stable isotopes
lightest generally in greater abundance
(C,H,O,N,S)

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11
Q

isotope properties

A

same protons and electrons = same chemical behaviour

physiochemical differences

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12
Q

isotope properties, same chemical behaviour

A

enter same chemical reactions
form same bonds
rare isotope can trace abundant isotope

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13
Q

lighter stable isotope

A

generally more abundant

not Li, B

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14
Q

Important stable isotopes

A

H, D, C, N, O, S, Cl

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15
Q

isotope properties, physiochemical differences

A 
lead to differences in distribution between phases
boiling pt
freezing pt
density
vapor pressure, etc.
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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

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17
Q

low atomic mass, isotope effects

A

H, He, C, N, O, S, Cl

exception - Fe isotopes fractionated by bacteria

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18
Q

large mass difference, isotope effects

A

∆m D-H = ca. 100%
∆m 13C-12C = 8.3%
∆m 18O-16O = 12.5%

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19
Q

why measure stable isotopes as ratios

A

utility- compare identical species/phases

measurement - measuring ratios increases precision

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20
Q

δ

A

‰
unitless
differences between sample and standard readings
not absolute isotope abundance

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21
Q

natural abundance standard

A

defined as δ = 0

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22
Q

Stable isotope ratio notations

A

δ^n X sample (‰) = [Rsample - Rstandard/ Standard] x1000

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23
Q

R

A

absolute abundance ratio
atom% ^n X / atom% ^m X
e.g. atom% 15N /atom % 14N

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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
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25
Q

hydrogen stable isotope standards

A

SMOW, VSMOW

standard mean ocean water

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26
Q

oxygen stable isotope standards

A

SMOW VSMOW

standard mean ocean water

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27
Q

carbon stable isotope standards

A

PDB VPDB
Pee Dee Belemnite
13C/12C = 0.0112372
18O/16O = 0.0020671

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28
Q

nitrogen stable isotope standards

A

atm N2

atmospheric nitrogen

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29
Q

sulphur stable isotope ratio

A

CDT

Canyon diablo triolite

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30
Q

δ13C atmospheric CO2

A

-8.2‰

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31
Q

δ13C plants, kerogen, coal

A

-8 - -55‰

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32
Q

𝛿 13C oil

A

-22 - -50‰

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33
Q

𝛿13C natural gas

A

-25 - -100‰

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34
Q

range in hydrogen isotopic variation

A

-700 - +200‰

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35
Q

range in carbon isotopic variation

A

-140 - +40‰

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36
Q

range in nitrogen isotopic variation

A

-60 - +50‰

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37
Q

range in oxygen isotopic variation

A

-30 - +30‰

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38
Q

range in sulphur isotopic variation

A

-50 - +40‰

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39
Q

Hydrogen isotope mass differences

A

99.8%

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40
Q

Carbon isotope mass differences

A

8.36%

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41
Q

nitrogen isotope mass differences

A

7.12%

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42
Q

isotopes are subjected to

A

Isotope effects

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43
Q

higher mass differences =

A

larger isotope effects

more strongly fractionated

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44
Q

Isotope effects

A

departure in isotope ratio from global average abundance due to physiochemical mechanisms

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45
Q

types of isotope effects

A

EIE - equilibrium isotope effects

KIE - kinetic isotope effects

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46
Q

Isotope fractionation

A

expression of isotope effects

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47
Q

Oxygen mass differences

A

∆m 16O/18O = 12.%

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48
Q

sulphur isotope mass differences

A

∆m 32S/34S = 6.24%

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49
Q

causes for isotope effects

A

chemical, physical properties of isotope
physiochemical properties of isotopes
isotopologues

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50
Q

chemical, physical properties of isotopes

A

arise from differences in atomic mass
reaction rates
diffusion rates
equilibrium constants

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51
Q

physiochemical properties of isotopes

A

result of quantum mechanical effects
energy of molecule restricted to discrete E levels
e.g. heat capacity, density, vapour pressure

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52
Q

isotopologues of isotopes

A

same molecules with different masses
different vibrational energies
e.g. H2, DH, D2

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53
Q

different isotopologues

A

different masses = different vibrational energies = different zero-point energies

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54
Q

ZPE

A

zero point energy

energy difference between minimum in potential energy curve and ground state energy (Eo)

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55
Q

Eo =

A

1/2 hv
h = plancks constant
v = vibrational frequency

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56
Q

ZPE, heavy isotopologue

A

lower ZPE than lighter isotopologue because v varies inversely with mass

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57
Q

isotopologue bonds

A

weaker in light isotopologues, easier to break, due to higher ZPE

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58
Q

difference in chemical properties of isotopologues can be evaluated in terms of

A

vibrational frequencies

transition states, in case of kinetic effects

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59
Q

energy well, isotopologues

A

heavier isotope lower in energy well- escapes less easily

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60
Q

different ZPE =

A

fractionation during chemical reactions via 2 processes

  • equilibrium processes
  • kinetic processes
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61
Q

equilibrium processes

A

rare isotopes do not partition equally between equilibrating species or between different phases of same species at equilibrium

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62
Q

kinetic processes

A

isotopologues react at different rates in non-reversible reactions

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63
Q

molecular diffusion rates

A

differ between isotopologues b/c velocity of molecule depends on Ek and inversely on mass
H2 diffuses slightly faster than DH

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64
Q

different diffusion rates =

A

isotope effects

isotope fractionation

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65
Q

EIE

A

generally reversible rxn’s or physical processes
governed by ZPE (QMs)
permits isotope exchange

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66
Q

KIE

A

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

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67
Q

consequences of mass differences on isotopes

A

heavier isotope molecules have lower mobility

heavier isotope molecules have higher bonding energy

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68
Q

kinetic energy of a molecule

A

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

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69
Q

internal energy of gas molecules due to

A

translational energies
rotational energies
vibrational energies

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70
Q

internal energy of liquid, solid

A

stretching

vibrational frequency of bonds

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71
Q

12C18O bond strength

A

higher mass = low vibrational frequency = stronger bond

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72
Q

heavier isotopes have

A

lower mobility

higher bond energy

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73
Q

heavier isotope species tend to be

A

concentrated in more dense/strongly bonded phase

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74
Q

isotope fractionation between two phases

A

decreases w/ increasing T

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75
Q

isotope fractionation factor

A

α = Ra/Rb

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76
Q

ε =

A

(α-1)*1000

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77
Q

α_A-B is

A

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

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78
Q

fractionation factors expressed as 10^3lnα because

A

close approximation to permit fractionation between materials (ε)
value nearly proportional to inverse of T (1/T) at low T (ºK)

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79
Q

lnα varies as

A

1/T in low T

1/T^2 in high T

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80
Q

KIE typically

A

unidirectional A–> B
incomplete, not all of A reacted to B
relatively rapid

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81
Q

KIE examples

A

Evaporation
diffusion
enzymatic fixation
e.g. CO2 from atmosphere – plant; restricted by stoma

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82
Q

KE

A

kinetic energy, KE = 0.5mv^2

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83
Q

diffusion ratio

A

inversely proportional to mass
Da/Db = mb/ma
e.g. 12C16O2 diffuses 1.1% faster than 13C16O2

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84
Q

reduced mass

A

µ = (mi*M/mi + M)

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85
Q

why reduced mass used

A

collisions and interactions lower true diffusive rate

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86
Q

reduced mass diffusion ratio

A

“true ratio”

D1/D2 = sqr. root (µ2/µ1)

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87
Q

Dissociation energy

A

heavier molecules have higher dissociation energy

therefore bonds harder to break

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88
Q

For reactions that have not reached equilibrium or completion

A

light isotope is preferentially in the product pool

heavy isotope is in the reactant pool

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89
Q

evaporation under 100% humidity

A

almost equivalent to evaporation under closed-system conditions

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90
Q

condensation described by

A 
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
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91
Q

Rvap of condensation will be different for Oxygen than H why and how

A

𝛿18O will be less negative than 𝛿D because of the mass differences
D much greater than H so isotope effects are greater

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92
Q

single-phase, open-system evaporation under equilibrium

A

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

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93
Q

two-phase, closed system evaporation

A

𝛿18O of water and vapour increase but much less than open system
𝛿18O instantaneous and cumulative vapour identical

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94
Q

isotopic ratios on land

A

dependent on distance of transport

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95
Q

what kind of process is evaporation

A

kinetic process

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96
Q

assumptions of evaporation being a kinetic process

A

rapid

vapour carried away

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97
Q

D

A

deuterium

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98
Q

fundamental to understanding isotope systematics of hydrologic cycle

A

knowledge of isotope effects associated with evap and condensation between air masses, reservoirs

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99
Q

condensation is what kind of process

A

equilibrium process

easier to deal with mathematically, depends on T alone

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100
Q

larger f in Arctic water or equatorial?

A

Arctic
higher temperature = lower f
higher temperature = lower alpha

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101
Q

isotope effects in open system clouds, assumptions

A

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

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102
Q

If cloud undergoes condensate loss under equilibrium conditions and no exchange with environment, change in isotope ratio of remaining vapour

A

described by Rayleigh Distillation equation for closed system
Rt/Ro = f^(α-1)
f = fraction of cloud vapour remaining
α = Rcondensate/Rvapour

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103
Q

Rayleigh fractionation curve of cloud

A

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

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104
Q

larger isotope effects in clouds why

A

colder T than land where evaporation occurred

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105
Q

latitudinal variation in precipitation

A

15ºN: 𝛿18O = -2, 𝛿D = -6
25ºN: 𝛿18O=-5, 𝛿D = -30
60ºN: 𝛿18O = -15, 𝛿D = -110

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106
Q

areas of similar rain composition on a map

A

isoisotopic lines

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107
Q

isotopes in rain controlled by

A

latitude (# of rain events)
altitude (T)
distance from coast (# of rain events)

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108
Q

GMWL

A 
global meteoric water line
𝛿Dsmow vs 𝛿18Osmow
𝛿D = 8𝛿18O + 10
y axis = 0 down to -500
x axis = -50 right to 0
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109
Q

Hydrogen isotope uses

A

hydrology
water
mineralogy
geothermometry

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110
Q

Helium isotope uses

A

mantle, subsurface geochemistry

pathway tracer

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111
Q

Carbon isotope uses

A

life
biology
partitioning or organic/inorganic compounds, pools
geothermometry

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112
Q

Nitrogen isotope uses

A

life
trophic levels
biological processes

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113
Q

exogenic carbon cycle

A

outside of Earths interior

recycled at surface

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114
Q

Major crustal carbon reservoirs

A

organic carbon (life)
continental crystalline rocks (graphite, diamond)
**sedimentary inorganic C / carbonates (limestones)

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115
Q

Amount of carbon in reservoirs

A 
atmosphere 800 PgC (10^15)
ocean ca. 35,000PgC
land plant ca. 1000PgC
sedimentary 5x10^22gC
Corg 15x10^21gC
crust 7x10^21
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116
Q

∂13C for C reservoirs

A

atmos: -8.3‰
ocean: TDC=0‰, DOC=-20‰
land plants:-25‰
sedimentary: 0-1‰
organic: -23‰
crust: -6‰

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117
Q

oxygen isotope uses

A 
paleogeoscience
hydrology
water
mineralogy
geothermometry
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118
Q

sulphur isotope uses

A

global and regional redox state

biological (bacterial) processes

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119
Q

strontium isotope uses

A

Earth history

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120
Q

∂13C surface ocean

A

1.8‰

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121
Q

∂13C deep ocean

A

0.6‰

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122
Q

∂13C fossil fuels

A

-28‰

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123
Q

why is there different isotope range for primary producers

A

source C: -8‰ diff. between marine/atmosphere

diff. fractionation processes w/i PP: C4/phtyo./C3

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124
Q

C3

A

Calvin-Benson
‘normal’ plants
cooler, wetter, cloudier climates

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125
Q

C4 plant

A

Hatch-Slack
evolved for low CO2 in Cenozoic (65Mya)
bright, dry, warm places
more efficient with water, less efficient with light

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126
Q

examples of C4 plants

A

maize
sugar cane
desert plants

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127
Q

‘background’ CO2

A

180ppm

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128
Q

plant minimum

A

80-100ppm

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129
Q

difference between C3, C4

A

C3 takes CO2 into mesophyll cell and directly into C-B cycle

C4: CO2– mesophyll– bundle sheath cell– C-B cycle

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130
Q

function of bundle sheath

A

concentrates CO2

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131
Q

CAM

A

Crassulacean Acid Metabolism plants

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132
Q

what are CAM plants

A

have C3 and C4 system that are used separately

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133
Q

what is biggest problem plants have

A

close stomata when dry conditions to eliminate evaporation - suffocate
C4 plants are advantaged in this way because concentrate CO2 - have some stores

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134
Q

how CAM systems work

A

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

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135
Q

where are C3 plants

A

ubiquitous- all aquatic and ancients

high latitudes, cool climates, forests, woodlands, high latitude grasses

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136
Q

∂13C C3

A

-23 - -33‰

average -26‰

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137
Q

C4 common plants

A

tropic/warm grasses, spartina (marsh plant)

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138
Q

when C4 is most favourable

A 
p(CO2) less than 500ppm
high p(CO2) C4 is less favourable than C3
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139
Q

∂13C C4

A

-9 - -16‰
average -13‰
higher plants (angiosperms) -10 - -18‰
10-14‰ more enriched in 13C than C3

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140
Q

∂13C CAM plants

A

-9 - -33‰

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141
Q

Isotopic range of petroleum

A

bimodal due to presence of C3/C4 plants, more from C3 (aquatic plants), higher in the -20 - -30‰

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142
Q

C3/C4 plant distribution on earth

A

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

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143
Q

∂13C coal

A

-25‰

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144
Q

∂13C natural gas

A

-41‰

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145
Q

∂13C petroleum

A

-28‰

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146
Q

∂13C anthropogenic CO2

A

-26‰

making atmosphere lighter, more negative

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147
Q

urban atmosphere ∂13C

A

-12‰

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148
Q

what happens to isotope ratios during burial and decomp

A

become heavier

C3 plants: -27‰ –burial - soil org matter -27‰ bacterial decay and respiration(+5‰) - soil CO2 -22‰ – (+10‰) — -12‰

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149
Q

observed discrimination (∂13Catm - ∂13Cplant) vs. (CO2)int/(CO2)ext

A

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

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150
Q

why is there observed discrimination

A

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

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151
Q

∂13C of photosynthesizers

A

algae: -10 - -22‰
plankton: -18 - -31‰
kelp: -10 - >-20‰

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152
Q

∂13C of photosynthesizes dependent on

A

pCO2, T, S, pH
source (atmosphere vs water)
cytoplasm ƒ

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153
Q

temperature effects on ∂13C of photosynth.

A

higher T = lower fractionation
colder water = more CO2
equatorial plants = lower fractionation
CO2 diffusion rates

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154
Q

land plant ∂13C

A

-32‰ - -22‰

155
Q

algae ∂13C

A

-22‰ - -10‰

156
Q

∂13C source

A

∂13C(HCO3) = 0‰ (-0.2‰ in water)
∂13C(CO2) = -8‰ in air
aquatic plants likely use HCO3 and CO2

157
Q

cytoplasm CO2

A

in aquatic environment CO2 same phase as cytoplasm CO2 - no isotope effect to convert it

158
Q

CO2 diffusion

A

slower in water than air

able to be used more efficiently in aquatic plants (similar to differences in C3 vs C4)

159
Q

marine vs freshwater plants

A

marine typically enriched
marine plankton ∂13Corg -28 - -17‰
fresh plankton ∂13C -20‰ - -32‰

160
Q

why difference in marine vs fresh

A

HCO3 source

161
Q

HCO3 ratio

A 
∂13C(HCO3)marine = 0‰
∂13C(HCO3)fresh = -9‰ - -15‰
162
Q

HCO3 source

A

ocean - mostly from atmosphere

lake - mostly from groundwater

163
Q

∂13C(HCO3)freshwater depends on

A

residence time
source
degree of equilibration

164
Q

∂13C(HCO3)freshwater residence time

A

mixing, exchange with atmospheric CO2

165
Q

∂13C(HCO3)freshwater source

A

water body size
circulation
temperature

166
Q

∂13C(HCO3)freshwater degree of equilibration

A

equilibration w/ ∂13C(HCO3)groundwater

167
Q

plankton c-isotope dependence on T

A

largest α at lowest T = strongest isotope effect = lightest/most negative at high/low latitudes

168
Q

plankton c-isotope dependence on T, why

A 
enzymatic T-dependence (lower alpha at higher T)
CO2 solubility (more soluble at lower T)
169
Q

why does high [CO2] affect plankton ∂13C

A

high [CO2] = higher diffusion supply to plankton = large fractionation (like C3 plants)

170
Q

C isotope dependence on trophic level

A

plants - herbivores - 1ºconsumer - 2ºconsumer
ratio becoming less negative as go –>
e.g. -25‰ -(-18‰) –(-16‰ ) –(-15/-10‰ )

171
Q

how to estimate T history of sea water and volume of ice caps

A

use long-term persisting records - fossil limestone (calcite)

172
Q

Harold Urey

A

first to O2 isotopes ratios of carbonates to deduce T of carbonate deposition, now cornerstone of paleooceanography

173
Q

Basic idea of carbonate paleothermometer

A

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
Q

obstacles to paleothermometry

A 
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
Q

∂18Ocarbonate formation, equilibrium

A

disequilibrium if there was rapid precipitation

176
Q

disequilibrium formation of ∂18O carbonate

A

biogenic vital effect

177
Q

example of biogenic vital effect

A

plankton bloom (growth outside of equilibrium)

178
Q

∂18Ocalcite change with time

A

post-burial isotopic exchange w/ pore water

dissolution, recrystallization, etc.

179
Q

paleothermometry methodology

A

liberate CO2 from CaCO3 by dissolution w/ H3PO4

180
Q

number of CO2 isotopologues

A

12

181
Q

most common CO2 isotopologues

A

12C16O16O
13C16O16O
12C18O16O

182
Q

CO2 liberation from CaCO3 equation

A

3CaCO3 + 2H3PO4 -> 3CO2 + 3H2O + Ca3(PO4)2

183
Q

12C16O16O

A

mass 44
most common isotopologue
98.450 mole%

184
Q

13C16O16O

A

mass 45

1.065 mole %

185
Q

12C18O16O

A

mass 46

0.405 mole %

186
Q

C/O isotopes in carbonates referred to

A

VPDB

187
Q

O isotopes in water referred to

A

VSMOW

188
Q

carbonate EIE

A

O2 in bicarbonate equilibrates isotopically with O2 in water
bicarbonate used by organisms w/ shells

189
Q

carbonate EIE equations

A

Ca2+ + 2HCO3- CaCO3 + H20 + CO2

H2O + CO2 2HCO3-

190
Q

bicarbonate isotope fractionation factor

A 
α_c-w = Rcalcite/Rwater 
R_c = 18O/16O in calcite
191
Q

K of oxygen isotope

A 
K = ([CaC18O3]/[CaC16O3])^1/3  / ([H218O]/[H216O])
K = Rc/Rw = α
192
Q

Oxygen isotope recorder equation

A

10000 lnα_c-w = 2.78(T^-2 x 10^6) - 3.39 (T in K)

193
Q

calcite-aragonite fractionation

A

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
Q

why use Calcite-Aragonite relationship to determine T

A

allows you to get around needing to know H2O characteristics
Arag water calc.
arag calcite

195
Q

calcite-aragonite offset at high T

A

converge

lower α = less discrimination

196
Q

∂18Ocarbonate rule of thumb

A

at constant ∂18Oseawater:

change in ∂18Ocarb of ca. 1% corresponds to ca. 4ºC change

197
Q

using ∂18Ocarb rule of thumb

A

use to find T of calcite precipitation only if ratio of water is known
∂18Ocalcite = ∂18Owater - 0.23∆T

198
Q

using foraminifera

A

planktonic = surface water T

benthic formas = deep water T = long term/ background/ 1000yr scale/ baseline

199
Q

benthic foraminifera

A

18O enriched (cold water)

200
Q

temperature record from belemnite shell growth rings

A

can see inter-annual seasonal variability

201
Q

comparing planktonic to benthic foraminifera

A

benthic = colder = heavier ∂18O
planktonic = warmer = light/depleted ∂18O
use offset to understand changes in T

202
Q

carbonate paleotemperature equation variables

A

∂18Ocarbonate
∂18Owater
Temperature

203
Q

how T is found for carbonate paleotemp. equation

A

estimate T from measured ∂18Ocarb assuming ∂18Owater value

204
Q

validity of T estimate in carbonate paleothermometer depends on

A

∂18Owater at time of calcite growth
∂18Ocarbonate alteration
carbonate precipitation equilibrium with water

205
Q

∂18Owater at time of calcite growth

A

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
Q

∂18Ocarbonate alteration

A

metamorphism

low T diagenesis

207
Q

how can we tell if there is ∂18Ocarb alteration

A

thin sections!

208
Q

carbonate precipitate in equilibrium with water?

A

vital effect

some organisms secrete carbonate not in equilibrium

209
Q

∂18Owater today

A

30±15‰

ice caps/glaciers = 2.5% of hydrosphere

210
Q

variation in ∂18Owater

A

ca. 1.2% between glacial max and interglacial

0. 1‰ / 10m

211
Q

when ice sheets are more expansive ∂18O

A

more positive than 0‰

212
Q

∂18Owater error of 0.2‰

A

corresponds to T error of 1ºC

213
Q

glacial ice can be estimated by

A

sea level changes

214
Q

sea level, glacial ice change used

A

make estimates of increase in ocean ∂18Owater during glacial times

215
Q

problem with using glacial ice change to estimate changes in ∂18Owater

A

does not consider sea ice (floating, no effect on sea level)

216
Q

sea ice ratio compared to water

A

higher 18O/16O ratio

217
Q

sea level at last glacial extent

A

-120m

218
Q

whats happening to ∂18Owater now

A

adding light water due to glacial melt

219
Q

foram isotope curves from caribbean cores

A

show regular changes in sea level associated w/ glacial/interglacial periods
tropical cores less affected by climate change

220
Q

high latitude snow

A

light
heavy 18O rich water condenses on mid-lats
atmos. water vapour increasingly depleted in H and O

221
Q

Interior Antarctica 18O

A

5% lower than ocean water - meltwater from glacial ice is 18O depleted

222
Q

interglacial sea level (today)

A

mean ocean depth = 4000m

∂18O = 0‰ (SMOW)

223
Q

glacial sea level

A

∆sea level = 120m
120m/4000m = 0.03 (3% of oceans water frozen)
∆∂18Owater = 1.2‰

224
Q

why is glacial period sea water ∂18O = 1.2‰

A

∆∂18O difference between ocean (0‰) and ice (-40‰)

0.0‰ - (-40‰ x 0.03) = 1.2‰

225
Q

why important to know ∂18O = 1.2‰ during glacial

A

organisms growing there will have very different isotopic signatures - must correct for the ∆

226
Q

how to correct for the ∆∂18O from glacial times

A

using benthic water

227
Q

∂18O rich carbonate =

A

colder T

228
Q

main paleoceanography question

A

hw much of ∂18O shift is due to ice volume (sea level change) and how much due to T change

229
Q

LGM

A

last glacial max

230
Q

LGM amount of water in ice caps/glaciers

A

6.5%

231
Q

∂18O changes with glacial advance on short time scales

A

ca. 100,000yr entire ocean mass increase ca. 1‰ as light water transferred to ice sheet
marine carbonates change in accordance

232
Q

what causes changes in carbonate signature during glacial extent/retreat

A

isotopic shift of water

change in water temperature

233
Q

Effect of T increase on 18O precipitation

A

increased T = lower ƒ_carb-wat

organisms precipitate carbonate w/ lower ∂18O values

234
Q

Effect of T increase on 18O precipitation is compounded by

A

melting ice caps = reintroduced light water

∂18O of carbonates will also decrease for this

235
Q

effects of changing T and seawater isotopic composition during glacial cycle

A

complimentary effects

cause ∂18O value of marine carbonates to change in same direction

236
Q

how to circumvent the complimentary T/isotope effect (glacial periods)

A

use tropical organisms (little-no T effect)

use benthic organisms (little-no T effect)

237
Q

cenozoic glacial-interglacial periods

A

13 stages over last 500ka

238
Q

heavier isotope species tend to

A

concentrate in more dense phase

bond more strongly

239
Q

isotope fractionation between two phases tends to

A

decrease with increasing temperature

240
Q

what happens to cloud as it loses moisture

A

depleted (more negative)

241
Q

geographic distribution of isotopes in water

A

lighter/depleted towards poles

hydrogen shows larger range due to mass differences (0- -270‰)

242
Q

Deuterium excess

A

d = ∂D - 8∂18O

243
Q

global meteoric water line

A 
∂D = 8∂18O + deuterium excess
∂D = 8∂18O + 10 
d = 10 for GMWL
244
Q

18O vs 16O mass difference

A

2amu

245
Q

2H vs 1H mass difference

A

1amu

246
Q

what is deuterium excess

A

reflects slower movement of H218O vs HDO during diffusion = enrichment of HDO in less strongly bound phase (gas if water evaporation)

247
Q

LEL

A

local evaporation line

slope less than 8 as in GMWL

248
Q

D-excess increase

A

in response to enhanced moisture recycle as a result of increased evaporation

249
Q

D-excess decrease

A

where water is lost by evaporation

250
Q

differences in d-excess from

A

varying T, relative humidity, sea surface wind speed

251
Q

d-excess, Canada

A

East coast: ca. 13 - 20

West coast: -4 - +4

252
Q

meteoric water in lebanon, isotopic composition determined by

A

3 main air masses:
Eastern Europe- humid, cold
Mediterranean sea - warm, rainy
Syrian desert - warm winds

253
Q

extra factors in lebanon MWL

A

mount induce high altitude isotope effect

254
Q

benefits of altitude effects

A

distinguish groundwater recharged at high altitudes vs low altitude recharge

255
Q

Eastern mediterranean d-excess

A

+22 b/c evap. processes at sea-surface occur due to low-humidity air masses of continental origin

256
Q

GMWL does not explain mediterranean so well, use

A

MMWL - mediterranean meteoric water line

257
Q

LMWL

A

Lebanese meteoric water line
different slope due to secondary evaporation during rainfall
∂D = 7.135∂18O +15.98

258
Q

air mass flow in mediterranean sea area

A

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
Q

∂18O vs altitude

A

∂18O (-9, -5)
altitude (0, 2500)
decreasing, linear
∆∂18O ca. 2‰ / km

260
Q

processes that shift values from MWL

A

O isotopes displaced due to exchange w/ volcanic CO2, limestone
H isotopes due to exchange w/ H2S, silicate hydration, clay dehydration

261
Q

MWL shifts in minewaters

A

strongly enriched in D, weakly in 18O
fall above MWL
form mixing lines w/ high 18O, 2H fluid

262
Q

why MWL shifts in minewaters

A

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
Q

isotope ratios in sediments are governed by

A

meteoric water signal
modulated by exchange, loss in subsurface
enriched by substitution w/ minerals in sediment

264
Q

additional isotopic impacts on sedimentary fluids

A

compaction further increases 18O, 2H

265
Q

connate fluid

A

formation fluid

266
Q

MWL shift based on T

A

shifts less at higher T because of reduced isotope effects

267
Q

Meteoric waters in oilfield brines

A

∂18O (0, -20)
degrees latitude north (30, 60)
decreasing ca. linear
AB nearly -20‰

268
Q

high T geochemistry isotope uses

A

geothermometry
reconstructing ancient hydrothermal system
detecting crustal assimilation in mantle-derived magma
tracing recycled crust in mantle
exploration, exploitation of geothermal resources

269
Q

gas isotope geothermometry

A

kinetics of CO2-H20 very fast
continuous re-equilibration
cannot determine Tmax
slowest rxn rate is CO2-CH4

270
Q

mineral isotope geothermometry

A

best have large T coefficient

eg. quartz-magnetite

271
Q

why quartz-feldspar is not an adequate mineral isotope geothermometer

A

∂18O T gradient is small

feldspar particularly susceptible to isotopic exchange w/ post-formational fluids

272
Q

stable isotope thermometry main principle

A

exchange

fractionation/partioning of heavy/light isotopes between coexisting phases

273
Q

most important role in isotopic fractionation

A

vibration motion

274
Q

fractionation between phases m and n defined as

A

10^3lnα_m-n = (α10^6 / T) + b (at low T)

275
Q

mineral - water oxygen isotope exchange

A

anhydrous mineral - water O2 exchange: b = 3.7

hydrous mineral - water O2 exchange = proportional to number of OH bonds in phase

276
Q

hydrous mineral - water O2 exchange

A

b = 3.7 * (1 - fraction of OH bonds / all O-bonds)
KAl3Si3O10(OH)1.8F0.2
b = 3.7 (1- (1.8/10))

277
Q

O2 isotope in gold exploration

A

18O depleted rock = most alteration due to pluton intrusion
groundwater percolation = hydrothermal mineral formation/deposition
O isotopes map alteration aureole

278
Q

geothermal gas fluids contain

A

CO2, CH4, H2, H2Ovapour

279
Q

a pair of geothermal gas fluids an

A

be used as geothermometer

280
Q

distribution of isotopes between components is a function

A

of temperature

281
Q

requirements of geothermometry

A

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
Q

how to avoid altering isotopic equilibrium when sampling

A

slow rate of isotopic exchange to prevent isotopic re-equilibration btw sampling/analysis

283
Q

well developed isotopic geothermometers

A 
CO2 + 13CH4 -- 13CO2 + CH4
CH3D + H20 -- HDO + CH4
HD + CH4 -- H2 + CH3D
HD + H2O -- H2 + HDO
CO2 + H218O -- CO18O + H2O
284
Q

in systems where gas species concentration is too low for isotopic T determination

A

SO4 + H218O – SO3*18O + H2O

32SO4 + H234S – 34SO + H232S

285
Q

why is there seasonal variation in the keeling curve

A

b/c plant activity lowers it every season

the top is the baseline

286
Q

∂13CO2 (atmosphere) vs year

A

decreasing
higher (more 13C) in summer due to higher 12C uptake by plants
lighter isotopes pumped into atmosphere (FF)

287
Q

seasonal oscillations in atmospheric ∂13C based on latitude

A

highest oscillation in N, decreases down to South Pole b/c of decrease in plant life - no drawdown effect

288
Q

atmospheric mixing

A

yr scale

289
Q

if atmospheric mixing is quick why is isotopic pattern maintained (variable)

A

because sources/sinks stay the same

290
Q

N-S atmospheric gradient

A

actually small even though most emissions are in NH

NH must be taking up large amounts of emissions

291
Q

average ∂13C of atmospheric CO2 over ocean

A

-7.5‰

292
Q

∂13C CO2 variation w/ time for NH and SH at mid-lats

A

-0.02‰ /yr

293
Q

∂13C CO2 continental air

A

-8 - 9‰

urban influence

294
Q

variations in ∂13C atmospheric CO2 over ocean

A

latitude
season
time

295
Q

∂13C CO2 continental air

A

large variation for variety of factors

eg uptake/release of biospheric CO2, combustion of ff

296
Q

1980 vs 2008 CO2

A

1980: 338ppm, -7.6‰
2008: 380ppm, -8.2‰

297
Q

Is atmospheric ∂13C change only from FF?

A

no, if it was we’d expect ∂13C to be -9.85

something else is impacting, e.g. deforestation

298
Q

∂13C atmospheric vs year

A

relatively stable at ca. 5 until 1800 where it suddenly and dramatically drops off

299
Q

the organic cycle can be represented by

A

CO2 + H2O — CH2O + O2

300
Q

∂13C global exogenic carbon reservoir

A

-5‰

surface earth average

301
Q

∂13C inorganic carbon reservoir

A

1‰

ƒ = 0.77

302
Q

∂13C organic carbon reservoir

A

average Corg = -26‰

ƒ = 0.23

303
Q

as organic carbon reservoir is depleted in 13C

A

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
Q

Corg/Ccarb

A

has been amazing constant over last 3.5Ga

amount of Corg buried, amount of PP relatively constant through time

305
Q

major extinctions in earth history

A 
Palaeozoic - Cambrian 550Ma
Ordovician 440 Ma
Devonian 354 Ma
P-T 250Ma
Triassic 195Ma
K-T 65
306
Q

great oxidation

A

2.3 Ga

no increase in Corg burial (none in any of the extinctions either)

307
Q

sulphur isotopes

A

32S (95.02%)
33S (0.75%)
34S (4.21%)
36S (0.02%)

308
Q

S standard

A

VCDT

Vienna Canyon Diablo Troilite

309
Q

general terrestrial ∂34S

A

-50 - +50‰

310
Q

Ocean ∂34S_SO4

A

+21‰

311
Q

∂34S_SO4 in past

A

Mesozoic: +10‰
Palaeozoic: +30‰

312
Q

exogenic S partitioned into

A 
inorganic S (gypsum, anhydrite)
organic S (anaerobic remineralization of OM)
313
Q

inorganic S

A

Ca2+ + SO4- + 2H2O – CaSO4*2H2O

alpha = 1.002 (small)

314
Q

organic S

A

remineralized by SRB
OM oxidation
Dissimilatory sulphate reduction
alpha = 1.025 (large

315
Q

SRB

A

sulphate reducing bacteria

316
Q

Organic matter oxidation

A

2CH2O + SO4- – H2S + 2HCO3-

317
Q

Dissimilatory sulphate reduction

A

SO4- + 2H+ + 4H2 —- H2S + 4H2O

318
Q

Sulphur isotope ratios on Earth

A

inorganics: igneous, volcanics, petroleum, coal = 0 - +40‰
organics: biogenic pyrite, shales, limestones = -50 - 0‰

319
Q

inorganic C/S isotopes governed by

A

EIEs

Keq = alpha_ A-B

320
Q

organic C/S phases governed by

A

KIEs
eg. Rayleigh eq
R_A = R_o*ƒ^alpha -1

321
Q

oxygenated S form

A

sulphate

322
Q

anaerobic S form

A

sulphide (pyrite)

323
Q

normal marine sediments C:S

A

8:3

324
Q

euxinic

A

anoxic

325
Q

euxinic sediment C:S

A

ca. 2:1

326
Q

non-marine sediment C:S

A

10:0.5

327
Q

Earth C/S Cycle

A

Basic Pool: OandA ∂34S=+19‰, ∂13C=1‰
Inorganic: ∂13C_carbonate=1‰, ∂34S_gypsum=19‰
Organic: ∂34S_shale=-16‰, ∂13C_shale= -26‰

328
Q

C/S enriched in organic phases

A

12C, 32S

329
Q

C/S enriched in inorganic phases

A

13C, 34S

330
Q

more burial of carbon implications for oxygen

A

more oxygen because back reaction does not occur
i.e. back rxn: O2 + CH2O — CO2 + H2O
CH2O buried, O2 remains

331
Q

more carbon burial implications for sulphur

A

more Corg burial = more net O2 = less pyrite (aerobic conditions)

332
Q

C/S ratio through time

A

increase at 400Ma

peak at 300Ma

333
Q

periods of high ∂34S

A

characterized by high rates of pyrite deposition

334
Q

increased burial Corg

A

higher ∂13C
higher atmospheric O2
oxidized sulphides to So4
lower ∂34S

335
Q

largest isotope effects in

A

KIE

especially enzymatic systems

336
Q

typical range in C isotopes

A 
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
Q

maceral

A 
coal building block
inertinite
vitrinite
exinite
sapropoels
338
Q

∂Dvs∂Corg for macerals

A

∂D (-130, -70)
∂Corp (-25, -23)
isotopic ranges w/i a single piece of coal

339
Q

periods of low ∂14S

A

likely had high rates of gypsum deposition

e.g. Permian

340
Q

why is there isotopic differences in coal

A

made up of plant pieces - different masses, diff water w/ diff isotope signatures, diff fractionation w/i plant

341
Q

hydrocarbon generation

A

cleave terminal position of long chain = methane = small, light molecule not strongly bonded

342
Q

isotope effect of cracking kerogen

A

mostly get 12CH4 - easiest to break off

343
Q

Petroleum formation

A 
T dependent
20-50ºC = bacterial CH4
50-150ºC = CO2, peak at 100
100-200ºC = thermogenic CH4
150-200ºC = H2S
344
Q

petroleum windown

A

80-200ºC

CO2, thermogen. CH4 dominant

345
Q

methane to atmosphere sources (increasing)

A

mantle gas, geothermal gas, water column, rice paddies, biogas, terrestrial hydrates, atmosphere, bacterial reservoirs, natural gas, coal/gas reservoirs, marine hydrates

346
Q

over methane sources

A

hydrothermal vents

347
Q

covert methane sources

A

wetlands

cows

348
Q

wetland methane emissions

A

ca 115 Tg CH4/yr

21% o total annual emissions

349
Q

cow emissions

A

ca. 55-60 kg CH4/cow/yr

350
Q

biogeochemical carbon cycling

A

CO2 – methanogenesis – CH4 – methanotrophy

351
Q

methanogenesis

A

carbonate reduction

epsilon_c = 49-95

352
Q

Methanotrophy

A

aerobic/anaerobic oxidation CH4 - CO2

353
Q

most reduced form of C

A

CH4

354
Q

Measure isotopic ratio, marine system

A

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
Q

Marine ∂D_CH4 =

A

∂D_H2O = 18O

356
Q

methyl fermentation ∂D_CH4

A

ƒ(∂D_H2O-18O) + (1-ƒ)*[(0.75∂D_methyl) + (0.25∂Hydrogen)]

357
Q

C-D diagram

A

∂13C_CH4 vs ∂D_CH4 = methane map, signatures of sources

Y shape - bacterial carbonate reduction, bacterial fermentation, geothermal/hydrothermal

358
Q

where does Atmospheric plot on C-D diagram

A

above and right of everything else

unique signature due to sink processes

359
Q

gas hydrates

A

10,000Gt C
10-1000X natural gas + coal + oil
meta-stable
will emit huge CH4 if released from permafrost melt

360
Q

Clathrate gun hypothesis

A

huge change in water mass signature 4-5‰- how to get that much ‘lightening’ - gas hydrate destabilization?

361
Q

Types of gas hydrate

A

Type I/II - biogas

Type H - thermogenic

362
Q

biogas hydrate ∂13C

A

∂13C -75 - -60‰

typically -65‰

363
Q

∂13C thermogenic

A

-35 - -55‰

typically -45‰

364
Q

hydrate ice worms ∂13Corg

A

sediments: -23 - -47‰

aragonite -40 - -45‰

365
Q

Methane diagenetic carbonate

A

isotopically very light
formed entirely of CH4 - methane oxidation
carbonate pavement precipitated from CH4
ugly brown, porous, inclusions

366
Q

carbonate ∂13C

A

diagenetic -20 - -25‰
methanogenic +10 - +26‰
methanotrophic -70 - -35‰

367
Q

hydrates and slumping

A

∆T - methane unfreezes - block of hydrate sediment breaks off - slumping/sliding /debris flow - gas plume released

368
Q

why do hydrate not usually release gas plumes

A

SRBs graze down methane - we rely on microbes to protect us

369
Q

disproving crath ray hypothesis

A

measure ice samples - see major [CH4] increase in younger dryas ca. 11.4kyr - measure ∂13C_CH4 - see no changes!

370
Q

what is the significance of no change in ∂14C_CH4 in younger dryas

A

if the [CH4] increase was due to methane hydrate release would expect a significant ‘lightening’ of the isotopic signature - do not!

371
Q

what could cause the increased [CH4]

A

no change in source, just more of it - wetland expansion?

372
Q

Snowball Earth

A

Cryogenian period, Neoproterozoic era, ca 650-700Ma- twice?

373
Q

how to get out of snowball earth

A

Volcanos! - erupting CO2, no life to take it up - accumulates = warming

374
Q

how to get in to snowball earth

A

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
Q

fate of CO2 mostly dictated by

A

weathering - ultimate CO2 sink

more weathering = more CO2 sink

376
Q

∂13C vs age, Meso, Neoproterozoic

A

leading in to Neoprot. see more highs/lows of ∂13C - transitioning in and out of snowball phases?

377
Q

∂13C_carb and burial

A

higher ∂13C_carb = more organic burial

378
Q

Alkenones

A

C37 ketones. di/tri-unsaturated long-chain alkenones

379
Q

using alkenones for paleothermometry

A

coccolithophores change amount of alkenone in membrane and therefore fluidity of membrane based on T

380
Q

Where do alkenones come from

A

uniquely derived by haptophytes (eg. coccolithophores)

381
Q

important coccolithophore

A

Emiliania huxleyi

382
Q

U^k_37 index

A

[C_37:2]/([C_37:3] + [C_37:3])

degree of unsaturation - fn of SST, specific to E Huxleyi growth

383
Q

why use alkenone-derive paleo barometry

A

can tell up to 30Ma, ice cores only tell ca 1Ma

2016/2017 (15 decks)

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