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EOS > EOS 260 Part II > Flashcards

EOS 260 Part II Flashcards

1
Q

steady state

A

dx/dt = 0
no change in position with time
stable or unstable

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2
Q

unstable steady state

A

unstable to a small perturbation- like on top of a hill
starting close to steady state system will diverge away
d2x/dt2>0

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3
Q

non-steady behaviour

A

transient

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4
Q

stable steady state

A

starting close to steady state system will converge to it
stable to a small perturbation
d2x/dt2<0

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5
Q

studying steady states

A

helps us understand the system

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6
Q

two stable steady states

A

must be separated by an unstable steady state

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7
Q

bistable system

A

chair, thermohaline circulation, geomagnetic reversals
can be resting in two states, states need not be symmetric with respect to stored energy
defining characteristic of bistability- 2 stable states (minima) are separated by a peak (maximum)

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8
Q

forcing

A

input to or control parameter affecting the system
will affect the state of the system
not affected itself by the state of the system

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9
Q

examples of forcing

A

solar forcing- sunlight earth receives

putting consecutively steeper ramps under a chair- force it to tip over

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10
Q

feedback

A

how the system will respond to a small perturbation or to a change in forcing
2 kinds

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11
Q

stabilizing feedback

A

negative feedback
diminishes effect of a perturbation, makes change smaller
pushes system back to stable steady state

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12
Q

Couplings

A

positive- change in A gives a change of same sign in B
shown by an arrow
negative- change in A gives a change of opposite sign in B
shown by line with dot

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13
Q

feedback loops

A

+1 to positive feeback
-1 to negative feedback
combined effect by multiplying them

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14
Q

most important gas absorbing insolation in atmosphere

A

H20 vapour

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15
Q

flux density units

A

W/m^2 = J / m^2 s

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16
Q

incoming solar radiation

A

341 W/m^2

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17
Q

destabilizing feedback

A

positive feedback
enhances the effect of a perturbation, makes change larger
pushes system on to next stable steady state after passing unstable steady state

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18
Q

outgoing longwave radiation

A

239 W/m^2

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19
Q

reflected solar radiation

A

102 W/m^2

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20
Q

total planetary albedo

A

outgoing/incoming

102/341 =~ 0.28 ~ 0.3

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21
Q

insolation absorbed by surface

A

161 W/m^2

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22
Q

average sunlight over whole earth

A

S/4 ~ 341 W/m^2

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23
Q

S

A

solar constant
S ~ 1368 W/m^2
relatively constant

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24
Q

where does insolation go

A

absorbed by atmosphere
absorbed by surface
reflected by surface
reflected by clouds/atmosphere

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25
Q

net energy absorbed by earth

A

0.9 W/m^2

absorbing ~1 W/m^2 more than it is emitting

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26
Q

where does surface radiation go

A 
thermals
evapotranspiration
emitted by atmosphere
back radiation- emitted back to surface
out of atmosphere- through atm. window
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27
Q

rate of temperature change

A

proportional to energy imbalance ∆F
inversely proport. to system mass x specific heat capacity
dT/dt = ∆F/mc
∆F x area (of earth)

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28
Q

sunlight in the atmosphere

A

reflected/deflected by clouds, atmospheric molecules

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29
Q

atmospheric molecules

A

why the sky is blue

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30
Q

area of earth

A

1.5x10^14 m^2

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31
Q

why did our dt calculation for 2k in the ocean underestimate the time it would take

A

we found 730yrs

our model assumes ocean mixing, which takes ~1000yrs

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32
Q

what can be seen from calculating dt of certain temperature change in the atmosphere vs. the ocean

A

ocean takes much longer to heat

global warming is dependent on the ocean

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33
Q

how many seconds in a year

A

60x60x24x365.25 = 31557600s

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34
Q

2 layers of the atmosphere

A

troposphere

stratosphere

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35
Q

troposphere

A

lower layer, T profile set by large scale convection, T decreasing with altitude

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36
Q

stratosphere

A

upper layer, minimal motion, T set by radiative absorption and emission (O3), increases with altitude

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37
Q

between troposphere and stratosphere

A

tropopause

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38
Q

before ozone was formed

A

stratosphere would be constant T with altitude

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39
Q

black body

A

absorbs all electromagnetic radiation- appears black

emits radiation dependent on wavelength and temperature

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40
Q

Planck function

A

describes radiative emission of a black body

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41
Q

Stefan-Boltzmann law

A

integration of the Planck function
F_BB = σ T^4
F is flux, σ is constant, T is in Kelvin

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42
Q

mean earth surface T and F

A 
T = 289 K
F_BB = 396 W/m^2
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43
Q

Wein’s displacement law

A

describes peak emission
as T of BB increases, emission peak moves to shorter λ
λ_max = b / T
λ_max is mx emission (µm), b is 2898µm K

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44
Q

some example λ_max

A

earth - 10µm (IR)
sun - ~500nm (visible)
largely different spectral regions

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45
Q

visible radiation absorption

A

not absorbed strongly by atmosphere

most radiation absorbed is at the surface

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46
Q

thermal radiation absorption

A

absorbed strongly by atmosphere

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47
Q

epsilon

A

emissivity band of thermal radiation absorption

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48
Q

Earth energy balance

A

sunlight absorbed

thermal radiation emitted

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49
Q

sunlight absorbed

A

earth absorbs energy from sun as a circle, some reflected back to space by clouds/atmos/surface

suns energy doesn’t reach all points of earth at one time

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50
Q

absorbing area

A

π r_e^2

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51
Q

emitting area

A

4πr_e^2

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52
Q

thermal radiation emitted

A

earth emits thermal energy from a spherical surface

emitting at all points on earths surface

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53
Q

effective temperature of emission

A

T_eff

if earth had no atmosphere/greenhouse effect

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54
Q

effective temperature equation

A

S/4 (1-alpha) = σ T_eff^4
S is solar constant = 1368 W/m^2
alpha is albed = 0.3

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55
Q

attenuation

A

is the gradual loss in intensity of any kind of flux through a medium

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56
Q

attenuation of radiation in atmosphere

A

Reflectivity (R), Absorptivity (A), Transmissivity (T)
R + A + T = 1
reflectivity is also called albedo

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57
Q

solar radiation in the atmosphere, attenuation assumptions

A

R = A = 0
T = 1
‘all insolation goes to E’s surface’

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58
Q

thermal radiation in the atmosphere, attenuation assumptions

A

R = 0
A ~ 0.75
T ~ 0.25
all IR is absorbed and transmitted, no albedo of IR

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59
Q

Kirchoff’s Law

A

A_λ = ε_λ
absorptivity = emissivity
only valid where the λ is the same on both sides

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60
Q

grey body

A

F_grey = ε σ T^4

some fraction of the planck function, a function of every λ

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61
Q

Radiative equilibrium equations

A

surface
S/4 (1 -alpha) + ε σ Ta^4 = σ Ts^4
atmosphere
ε σ Ts^4 = 2 ε σ Ta^4

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62
Q

rearranging equilibrium equations gives you

A

Ts^4 = 2Ta^4 surface T is always bigger

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63
Q

a greenhouse gas is

A

a gas that absorbs thermal radiation

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64
Q

2 most important greenhouse gases

A

CO2, H20

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65
Q

strength of the greenhouse effect felt by a gas

A

depends on logarithm of abundance

recall that an increase in 1 is an increased factor of 10 with logarithm

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66
Q

radiative forcing

A

change in net flux at tropopause (down minus up)

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67
Q

surface temperature changes is proportional to

A

radiative forcing

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68
Q

Earths surface T without greenhouse effect

A

~255K

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69
Q

greenhouse effect requires

A

absorption and re-radiation of thermal energy in atmosphere by greenhouse gases
that atmosphere is colder than surface

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70
Q

strongest radiative forcing change

A

if you add CH4 you’ll see a stronger effect than CO2 because the base value is so low, log properties, larger change

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71
Q

Climate feedbacks

A

Planck (temperature) feedback (-)
Water vapour feedback (+)
Ice-Albedo feedback (+)

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72
Q

Planck feedback

A

solar forcing–> Temp—>outgoing thermal rad—-.temp

increase T = more energy emitted = cools down

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73
Q

water vapour feedback

A

amplifies T change
solar F–>T—>Atmos. H2O vapour—>greenhouse effect—. outgoing thermal radiation
could lead to runaway greenhouse if Planck feedback stopped working

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74
Q

saturation vapour pressure

A

e_s is proportional to exp T
equilibrium between liquid and gas
increase water vapour = increase T
e_s is pH2O

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75
Q

greenhouse forcing depend on

A

change of logarithm of gas abundance

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76
Q

strength of water vapour forcing depends

A

linearly on temperature change

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77
Q

Ice-albedo feedback

A

solar F–>solar radiation absorbed—>T—.albedo—.solar radiation absorbed
cold–snow—higher albedo—colder
where there is snow we can see a lower temperature

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78
Q

albedo examples

A 
fresh snow 70-90
sea ice 50-75
desert 25-40
Forest/Grass 10-25
Ocean <10 -70 depending on incident ray
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79
Q

quick changing feedback

A

ice-albedo

there is snow, or there isnt

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80
Q

tau

A

thermal optical depth
strength of greenhouse effect
tau = tau_CO2 + tau_H2O

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81
Q

steady state energy balance on daisy world (no atmosphere)

A

σ T^4 = (1 - alpha)F_s

global albedo is a weighted sum of bare ground, black, white

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82
Q

solar flux on daisyworld

A

`increases linearly with time

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83
Q

daisyworld albedos

A

bareground: alpha_bare = 0.5
black: alpha_b = 0.25
white: alpha_w = 0.75

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84
Q

daisies are mesophile

A

growth rate beat depends on local T

max at 22.5ºC, no growth above 40ºC or below 5ºC

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85
Q

daisy growth

A

(A)(x)(beta)
x = p - A_w - A_b
p is proportion of planet surface that is fertile
A_w is area covered by white daisies

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86
Q

daisy death

A

gamma A

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87
Q

daisyworld model

A 
dA_w/dt = A_w ( x beta - gamma)
dA_b/dt = A_b ( x beta - gamma)
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88
Q

statement of Gaia hypothesis

A

Organisms and their material environment evolve as a single coupled system, from which emerges the sustained self regulation of climate and chemistry and habitable state for whatever is the current biota

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89
Q

what does Gaia hypothesis mean

A

life has controlling influence on physical/chemical climate
life changes atmosphere, maintains habitability through time
theory led to Earth System science

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90
Q

why earth is special

A

temperature and pressure allow all three phases of water

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91
Q

triple point of phase diagram

A

all 3 phases exist in equilibrium

pressure on y axis, temperature on x

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92
Q

cryosphere

A

the frozen water part of the Earth system

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93
Q

southern hemisphere snow

A

snowy year round, not a lot of year round change

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94
Q

northern hemisphere snow

A

highly variable, huge variation w/ season, ice albedo changes are a function of the NH

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95
Q

antarctic sea ice cycle

A

not a lot of variation due to Antarctic circumpolar current

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96
Q

arctic sea ice minimum

A

september

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97
Q

forming a glacier requires

A

accumulation of multi-yr ice on land

precipitation (snow) in winter and failure of this to melt fully in summer

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98
Q

consequence of temperature change with altitude in regards to snow/ice

A

mountains receive most precipitation
snowmelt least likely at high altitude, and on poleward facing slopes
mountain glaciers form first

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99
Q

glacier

A

valleys, follow topography

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100
Q

ice field

A

big glacier, still constricted by topography but ‘drapes’ topography, flow direction directed by topography

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101
Q

ice sheet

A

unrestricted by topography, covers it and goes where it wants

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102
Q

ice shelf

A

floats at edge of ice sheet/glacier

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103
Q

net ice build up

A

accumulation

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104
Q

net ice loss

A

ablation

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105
Q

above glacier equilibrium line

A

snow > melt

accumulation zone

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106
Q

below glacier equilibrium line

A

melt > snow

ablation zone

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107
Q

mountain glaciers in the absence of melt

A

can grow to form ice sheets

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108
Q

ice sheet-altitude feedback

A

growth of ice sheet–raises altitude of ice surface– surface temperature is colder– melt inhibited
positive feedback

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109
Q

ice sheet flow

A

from centre of ice dome (high pt.) outward

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110
Q

thicker ice sheet

A

more likely to slide at base

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111
Q

movement of ice as a function of base T

A

base > 0º, melt, wet base, glacier can slide, flow quickly, lubricated by melt-water percolated to bottom
base < 0º, bed frozen, glaciers flow more slowly, movement requires deformation

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112
Q

warming ice sheet

A

speed up flow–loss of altitude–warmer T–more melt

positive feedback

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113
Q

glacial records

A

glacial valleys, moraines, striations, tillites, loess

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114
Q

ice rafted debris

A

piece breaks off of iceberg–flows away– starts melting– drops sediment– can deform bottom sediment

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115
Q

glacial loess

A

When glaciers grind rocks to a fine powder, loess can form. Streams carry the powder to the end of the glacier. This sediment becomes loess

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116
Q

large ice cap instability

A

if ice reaches mid latitudes positive feedback will lead to global glaciation
from O-D energy balance model with ice-albedo feedback

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117
Q

small ice cap instability

A

if ice decreases too much it will all melt, no glaciation

ice sheets don’t get smaller, they collapse

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118
Q

north american ice sheets

A

lost: Laurentide, Cordillian, Scandanavian
remain: greenland

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119
Q

Phanerozoic/Proterozoic glaciations

A 
Oligocene-Present (poles)
Devonian, Cretaceous, Permian
Ordovician- Siluria
Cryogenian (Neoproterozoic)
Siderian (paleoproterozoic)
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120
Q

faint young sun paradox

A

with lower S glaciation would be deep (with todays atmospheric composition)
yet geological evidence shows less glaciation earlier in history
stronger greenhouse?

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121
Q

isotope

A

same number of protons, different number of neutrons, same atomic number, different mass, same chemical properties, different physical properties

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122
Q

oxygen isotopes

A

16O - 99.76%
17O - 0.04%
18O - 0.2%
16,18 most commonly used for palaeoclimate

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123
Q

hydrogen isotopes

A

1H - 99.984%

2H - 0.016%

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124
Q

2H

A

deuterium ‘D’

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125
Q

isotopologues

A

different forms of the same molecules, varying by isotopic composition
molecules with different isotopes, multiple heavy isotopes in one molecule, rare
ex. H2O with 0,1,2, deuterium

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126
Q

most common water isotopologues

A

(1H)2(16O)
(1H)2(18O)
(1H)(2H)(16O) - HDO
in order of commonality

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127
Q

isotopic ratio

A

R = rare X / common X

ex. D: R = D/H = 0.016/99.984 = 0.00016

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128
Q

problem with isotope ratio

A

variations are very small numbers

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129
Q

delta notation ratio

A

δX = 1000 ((Rsample-Rstandard)/Rstandard)

units are ‘permil’ parts per thousand

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130
Q

δD =

A

1000 [ ( (D/H)sample - (D/H)smow ) / (D/H)smow ]

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131
Q

SMOW

A

standard mean ocean water

standard for H and O for ice sheets

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132
Q

isotopic mass balance

A

m_tot = m_1 + m_2

m_tot δ_tot = m_1 δ_1 + m_2 δ_2

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133
Q

δ18Osmow =

A

0%o

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134
Q

negative delta isotope values

A

depleted in 18O
preferentially left behind
the more negative, the colder

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135
Q

saturation vapour pressure

A

describes amount of water vapour in equilibrium with liquid water

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136
Q

saturation vapour pressure of isotopologues

A

lower for heavier isotop., light isotopes evaporate preferentially
e_s((1H)2(16O)) is 1% lower than e_s((1H)2(18O))
e_s(HDO) is 10% lower than e_s((1H)2(16O)

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137
Q

water vapour isotopes compared to ocean water

A

water vapour is isotopically light
δ18O_wv < δ18O_ocean
δD_wv < δD_ocean

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138
Q

precipitation is isotopically heavy relative to water vapour

A

δ18O_wv < δ18O_precip

δD_wv < δD_precip

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139
Q

as temperature decreases from the source region of moisture

A

increasingly higher fractions of the atmospheric water vapour have precipitated out
remaining water becomes increasingly isotopically light and further precipitation will be lighter

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140
Q

isotopic concentration of precipitate a function of

A

temperature at which precipitation occurs

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141
Q

precipitation is isotopically lighter

A

further from the source

at higher altitudes

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142
Q

why ice cores are drilled at the summit of an ice cap

A

reduces distortion due to ice flow

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143
Q

stratigraphy of ice core with depth

A

~54m down- hard to see layers, use conductivity based on dust
~1800m down- clear layers can be seen
~3000m down- lots of sediment, layers may be lost
ability to resolve layers decreases with depth

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144
Q

ways to date stratigraphic layers

A

δ18O variation with seasonal cycle due to T variation
microparticle/glaciergeochemistry seasonal cycles
electrical conductivity- seasonal variations of contaminant load
reference horizons- radioactives, volcanic ash

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145
Q

trade off between time resolution of ice core and length of record

A

higher snow accumulation rate = better time resolution
faster accumulation = faster flow, shorter time record
longer record, lower resolution ex.greenland

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146
Q

why layers are thinner down ice core

A

pressure

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147
Q

why might δD be better than δ18O in ice

A

mass differences
H 1/2 (half of mass)
O 16/18

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148
Q

δD versus age graph

A

less (-), warmer T’s, interglacials

glacials last longer than interglacials

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149
Q

packed snow

A

firn

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150
Q

bubble formation in glacier

A

pressure of firn compresses packed snow into ice ~50m
100’s of years at warm, high accumulation
1000’s of yrs at cold, interior, low accumulation

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151
Q

age of air in ice

A

younger than ice itself (because they take so long to close)

and they close at different times

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152
Q

pacemaker of ice-ages

A

insolation changes (which affect ice volumes)

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153
Q

isotopes of ice sheets

A

accumulation of ice sheet removes light isotopes

ocean water become isotopically heavy

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154
Q

foraminifera

A

CaCO3 tests, must determine benthic vs. planktic

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155
Q

how forams represent ice volume

A

18O concentrated in CaCO3 relative to water (more with colder water), bust be planktonic (benthic waters don’t show much T variation)

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156
Q

marine δ18O graph

A

y-axis is opposite- heavy = high ice volume = cold

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157
Q

Pleistocen glaciation

A

regular glacial-interglacial cycles
before 700ka periodicity = 40ka
after 700ka periodicity = 100ka

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158
Q

change in glacial-interglacial periodicity

A

mid-Pleistocene climate transition

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159
Q

why are glacial studies better from antarctic

A

right on the pole-harder to melt ice sheet, more information contained in that
arctic ice is away from the pole, higher insolation

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160
Q

Croll-Milankovitch theory

A

Eccentricity
Axial tilt
Precession

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161
Q

eccentricity

A

how elliptical earths orbit it

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162
Q

eccentricity periodicity

A

100,000yrs

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163
Q

what eccentricity does

A

changes distribution of solar energy throughout year

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164
Q

perihelion

A

closest to sun (right now in NH Jan)

earth moving faster at this point of orbit to cover same amount of area

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165
Q

axial tilt

A

obliquity, the degree to which the axis is tilted changes

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166
Q

axial tilt periodicity

A

41,000yrs

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167
Q

axial tilt implications

A

changes strength of season

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168
Q

aphelion

A

earth farthest from sun

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169
Q

if obliquity were 0

A

no seasons

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170
Q

precession

A

wobble, change in where N/S axis points

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171
Q

precession periodicity

A

23,000 yrs

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172
Q

precession implications

A

relative strength and length of seasons

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173
Q

Croll-Milankovitch theory explains timing of glaciations

A

NH summer insolation is critical in determining ice volume, summer insolation determines whether glaciers survive ice-albedo feedback, most land mass- largest changes in ice extent

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174
Q

organic carbon dominantly fixed by

A

oxygenic photosynthesis
CO2 + H2O + hv —- CH2O + O2
inorganic matter reduced to organic matter

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175
Q

aerobic respiration

A

CH2O + O2 — CO2 + H2O

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176
Q

photosynthesis + aerobic respiration

A

closed cycle

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177
Q

decomposition of organic carbon in oxygen low environment

A

fermentation followed by methanogenesis

net: 2CH2O — CO2 + CH4

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178
Q

inorganic carbon

A

HCO3 -
CO3 2-
H2CO3

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179
Q

counter reaction to fermentation/methanogenesis

A

CH4 + 2O2 – CO2 + 2H2O

again, closed system

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180
Q

carbon reservoir cycling

A

biological reservoirs cycle quick

geochemical reservoirs cycle slow

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181
Q

atmospheric CO2 reservoir

A

760
photosynthesis takes 60 out
60 goes back in, fast cycling

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182
Q

residence time

A

tau = steady state reservoir size / flux

ex. 760 / 60 = ~13 yrs

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183
Q

basic structure of the ocean

A

2 layers: Mixed layer, deep ocean

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184
Q

mixed layer

A

~100m, geochemical equilibrium with atmosphere, well mixed by wind

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185
Q

ocean gradients

A

thermocline
chemocline
light penetration

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186
Q

thermocline

A

sudden small T decrease then, decreases gradually with depth down to ~100m where it reaches the minimum and stays ~constant

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187
Q

chemocline

A

ex. nutrient availability, abrupt decrease with depth

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188
Q

light penetration gradient

A

gradual decrease, area light penetrates = photic zone

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189
Q

Henry’s law

A 
assuming air-sea surface gas exchange in equilibrium
[x] = k_H px
[x] dissolved molar concentration (M)
k_H henrys law constant (M/atm)
px partial pressure (atm)
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190
Q

1atm =

A

101325 pa

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191
Q

gases are more soluble in

A

cold water

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192
Q

DIC reservoirs, smallest to biggest

A

H2CO3
CO2
CO3 2-
HCO3 -

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193
Q

carbonate chemistry

A

collective chemistry of DIC species

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194
Q

conserved oceanic carbon quantities

A

DIC

Alkalinity

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195
Q

DIC =

A

[CO2] + [H2CO3] + [HCO3 -] + [CO3 2-]

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196
Q

Alkalinity

A

expresses charge balance, net + charge from conservative ions balanced by net - charge of weak acids

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197
Q

circumneutral conditions in ocean

A

pH > 4

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198
Q

pH =

A
  • log [H+]
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199
Q

conservative ions

A

ions that don’t do much chemistry at ciarcumneutral conditions in ocean, salts, net +
[Na+] - [Cl-] + 2[Mg 2+] + 2[Ca 2+] > 0

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200
Q

Alkalinity, Alk =

A

[HCO3 -] + 2[CO3 2-] - [H+]

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201
Q

DIC reactions

A

CO2 + H20 ——– H2CO3
H2CO3 ——– H+ + HCO3 -
HCO3 - ——– H+ + CO3 2-

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202
Q

increasing pCO2

A

increases [CO2] – increases [H2CO3] — increases [H+] and [HCO3-] — increases [HCO3-] —- decreases [CO3 2-]
increased hydrogen ions, decreases pH, ocean acidification

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203
Q

why do we subract [Cl-] in conservative ion equation

A

because we are adding up the positive charges

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204
Q

pCO2 in DIC vs. Alkalinity graph

A

below ~where we are now (~0) - not possible

increasing in DIC = increased pCO2 (up to ~6µatm)

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205
Q

CO3 ^2- in DIC vs. Alkalinity graph

A

proportion of CO3 that is DIC
now ~0.2, slightly above ‘possible range’
increases in a very narrow band with increasing alkalinity up to 1
narrow band beyond the 1 that is 1 - (~0.2)
majority of the graph space is 0

206
Q

HCO3 ^- in DIC vs Alkalinity graph

A

proportion of DIC that is HCO3 ^-
currently pretty close to 1
thin strip going down to zero below where we are now, most of graph space is above where we are now and decreases gradiently to 0

207
Q

CO2 in DIC vs. alkalinity graph

A

proportion of DIC that is CO2
currently 0
increases gradiently above present value, most of graph space is 1

208
Q

pH in DIC vs. alkalinity graph

A

currently 7-9
7+ is all in a very narrow band, around where we are now
above present day values graph space is 5-3

209
Q

where we are now in terms of DIC and alkalinity

A

DIC ~ Alkalinity ~ 10^3

210
Q

If DIC were&raquo_space; alkalinity

A

pCO2 - high (6µatm)
pH - low (3)
~all Carbon is in CO2

211
Q

if alkalinity&raquo_space; DIC

A

can’t be

this is the ‘impossible’ part of the graph

212
Q

if alkalinity and DIC both maxed

A

pCO2 medium (~3-4)
pH - ~8
most Carbon in HCO3 ^1
note that this is very similar to current conditions- staying along line of same slope

213
Q

why - [H+] in Alk equation

A

adding up the negative ions

** H+ is generally negligible anyway

214
Q

max alkalinity

A

all carbon in carbonate
Alk = [HCO3] + 2[CO3] - [H+]
the 2 coefficient is key, how to maximize equation

215
Q

if Alkalinity = DIC

A

all carbon in HCO3

216
Q

alkalinity high relative to DIC

A

carbon mostly in carbonate

very little carbon in CO2

217
Q

decrease in alkalinity : DIC

A

transition to CO2 dominance

218
Q

weathering rocks

A

adds to DIC and alkalinity

219
Q

adding CO2 from atmosphere

A

only increases DIC

220
Q

alkalinity units

A

we have been saying M (concentrations), but its actually supposed to be measured in equivelants/L

221
Q

solid calcium carbonate precipitation from ocean

A

Ca^2+ + CO3^2- ——- CaCO3(s)

222
Q

solubility product

A

k_sp = [Ca^2+]_sat . [CO3^2-]_sat
products over reactants, only reactant is solid {1}
note that this equation is derived from the reverse equation of CaCO3 precipitation

223
Q

saturation product

A

Ω = [Ca^2+].[CO3^2-] / k_sp

224
Q

Ω < 1

A

subsaturated

CaCO3 dissolves

225
Q

Thermodynamic Theory

A

Ω < 1 subsaturated
Ω = 1 saturated
Ω > 1 supersaturated
Ω > 3 biological precipitation

226
Q

Ω > 1

A

supersaturated

CaCO3 precipitates

227
Q

Ω > 30

A

authigenic precipitation (abiotic)

228
Q

CaCO3 polymorphs

A

calcite

aragonite

229
Q

aragonite properties

A

k_sp: 6.5x10^-7 mol^2/kg^2

Ω: 4

230
Q

calcite properties

A

k_sp: 4.3x10^-7

Ω: 6

231
Q

more soluble CaCO3 polymorph

A

aragonite- lower Ω, get to saturation point at lower concentration

232
Q

decrease in ocean [CO3 ^2-]

A

decrease in saturation product (Ω)

problem for calcifying organisms

233
Q

k_sp,cal/arag =

A

( [Ca2+].[CO3^2-] )_sat,cal

234
Q

Ω_cal/arag =

A

[Ca2+][CO3^2-] / k_sp,cal/arag

235
Q

variations in k_sp dominated by

A

depth (pressure)

236
Q

as pressure increases k_sp

A

increases (solubility increases)

237
Q

solubility increase with depth

A

2X / 4km depth

238
Q

biologically precipitated CaCO3

A

sinks when organisms die– sedimented or dissolved in deep

239
Q

CCD

A

carbonate compensation depth
rain rate of CaCO3 = CaCO3 dissolution
no sediment below

240
Q

preservation of CaCO3 in sediment

A

~20-30%

241
Q

depth that Ω = 1

A

saturation horizon, Lysocline

CaCO3 starts to dissolve

242
Q

lysocline graph

A

Ω relatively high in surface
decreasing to saturation horizon
stays at 1 with increasing depth until reaches CCD then decreases further to 0

243
Q

change in [CO3^2-] dominate

A

over changes in [Ca2+]

244
Q

surface [CO3^2-] =~

A

250µmol/kg

245
Q

Atlantic depth characteristics

A

deep ocean [CO3^2-] - 113µmol/kg

saturation horizon - 4.5km

246
Q

Indian ocean depth characteristics

A

deep ocean [CO3^2-] - 83µmol/kg

saturation horizon - 3km

247
Q

North Pacific ocean depth characteristics

A

deep ocean [CO3^2-] - 70µmol/kg

saturation horizon - <1km (not much CaCO3 sediment)

248
Q

more deep ocean [CO3^2-]

A

deeper saturation horizon

Ω = Ca x CO3 / ksp — increase CO3 = increase Ω

249
Q

deepest saturation horizon of the 3 oceans

A

Atlantic

250
Q

changes in surface when CaCO3 precipitates

A

alkalinity decreases by 2 units

DIC decreases by 1 unit

251
Q

changes in deep ocean when CaCO3 dissolves

A

Alkalinity increases by 2 units

DIC increases by 1 unit

252
Q

carbonate pump

A

moves CaCO3 from surface to depth

253
Q

why ocean carbon reservoir is so large

A

more C is stored in deep ocean than there would be if ocean was in equilibrium with atmosphere

254
Q

Walker 1983

A

carbonate sedimentation throughout earths history constrains oceans chemistry and atmospheric composition

255
Q

DIC is fixed to organic carbon by

A

photosynthesis in surface ocean

256
Q

photosynthesis

A

CO2 + H2O == CH2O + O2

257
Q

total organic carbon

A

GPP

258
Q

GPP

A

gross primary productivity

259
Q

total organic carbon after respiration by producers

A

NPP

260
Q

NPP

A

net primary productivity

261
Q

carbon sinking out of photic zone due to organism death

A

export production

262
Q

in deep ocean most sinking carbon is

A

respired

263
Q

carbon in deep ocean that is buried in sediment

A

burial flux

264
Q

burial flux percentage of carbon in deep ocean

A

0.1%

265
Q

organic carbon buried in sediment

A

massive deposits

266
Q

massive deposits

A

black shale

up to 10% organic carbon

267
Q

how much organic carbon is needed to turn sediment black

A

only a few %

268
Q

pockets of organic carbon

A

diffuse deposits, kerogen

269
Q

organic carbon fluxes

A

GPP - 100GtC/yr
NPP - 50 GtC/yr
export production - 10 GtC/yr
Burial flux - 0.1 GtC/yr

270
Q

organic carbon that makes it out of surface ocean into deep ocean

A

~10%

271
Q

carbon isotope ratios

A

12C - 0.9893
13C - 0.0107
14C - 10^ -17

272
Q

stable carbon isotopes

A

12C

13C

273
Q

molecules with different isotopic compositions

A

isotopologue

ex. 13CO2

274
Q

δ (‰) =

A

[ R_sample / R_standard - 1 ] x10^3

275
Q

R_sample or R_standard =

A

heavy isotope / light isotope

ie. 13C/12C

276
Q

carbon standard

A

Pee Dee Belemnite (PeeDee formation)

277
Q

R [13C/12C] PDB

A

0.011237

δ 13C= 0‰

278
Q

R [13C/12C] CO2_atmospheric

A

0.01114486

lower than ocean

279
Q

( - ) isotope values

A

depleted

280
Q

permil

A

no units

≠ ppt (mass/volume)

281
Q

δ 13C _CO2, atmospheric

A

-8.2‰

282
Q

δ 13C_organic

A

~ -25‰

-20 - -40‰

283
Q

δ 13C_carbonates

A

~ 0‰

284
Q

δ 13C_HCO3-

A

~ 0‰

285
Q

δ 13C_volcano

A

-5‰

286
Q

δ 13C_fossil fuel

A

very low

graph of δ13C vs time is a decreasing trend, almost exact opposite of CO2 accumulations

287
Q

fractionate

A

vibrational movements/energy differs btw. isotopes
light isotope bonds are weaker relative to heavy
light isotopes are easier to evaporate relative to heavy

288
Q

vibrational movement graph

A 
# of atoms vs. velocity
normal distribution, most atoms in mean area, some in tails
tails of distribution are longer in light isotopes
289
Q

fractionation steps in plants

A

diffusion

carboxylation

290
Q

Diffusion

A

ε = 4.4‰
lighter molecule enters cell preferentially
cell depleted in 13C

291
Q

carboxylation

A

ε = (-15) - (-30)‰
energetically expensive process
easier to break bond in light isotopes

292
Q

ε

A

difference in δ values between 2 species

293
Q

after 2 steps of fractionation

A

cell becomes ~ -25‰

but this depends on whether we are talking about land or ocean plants

294
Q

why does overall plant cell depletion depend on land/ocean

A

source of original carbon that is taken into cell in step 1
air: -8.2‰
HCO3- : 0‰

295
Q

surface carbon cycle

A

F_in = F_org + F_carb = F_out

296
Q

F_in

A

volcanoes

-5‰

297
Q

F_org

A

ΔB

-25‰

298
Q

F_carb

A

δ 13HCO^3-

0‰

299
Q

ΔB

A

biologic fractionation

300
Q

f_org =

A

F_org / F_in
= δin - δcarb / δorg - δcarb
= ( -5 - 0 ) / ( -25 - 0 ‰ ) = 1/5 = 0.2

301
Q

which F is stable through earths history

A

Fin

302
Q

% of carbon buried as organic carbon

A

20%- preserves oxygenated state of atmosphere

303
Q

increase in δ 13C_carb

A

more f_org

higher productivity

304
Q

increase in f_org

A

burying alot of light carbon so δ13C increases because f_carb decreases, have to maintain isotopic balance

305
Q

represent ordinary continental silicate rocks as

A

CaSiO3

306
Q

ordinary continental silicate rocks

A

granite, basalt

307
Q

carbonic acid in rain

A

2[CO2 + H2O — H2CO3]

308
Q

continental rocks weathered by carbonic acid in rain

A

CaSiO3 + 2H2CO3 — Ca^2+ + 2HCO^3- + SiO2 + H2O

309
Q

consequence of weathering continental rocks

A

draws down 2CO2 from atmosphere, transfers Ca,2HCO3 to ocean

310
Q

continental rock weathering drawing down CO2

A

2CO2 removed from atmosphere = 2 units of DIC

311
Q

continental rock weathering transferred ions

A

add 2 units of DIC

add 2 units of Alk

312
Q

net changes from weathering of continental rock

A

ocean gain:
0 units DIC
2 units Alk
atmos. CO2 cancelled out by ocean CO2 (Henrys Law)

313
Q

why can we call the atmospheric CO2 ~ ocean CO2, and cancel out gains and losses?

A

because they equilibrate so rapidly

314
Q

continental rock weathering shifts carbon

A

from CO2 – CO3 ^ 2-

315
Q

shifting carbon from CO2 to CO3 ^2- results in

A

increase in CaCO3 saturation

higher CaCO3 deposition

316
Q

weathering continental rock, shifting carbon equations

A

Ca^2+ + 2HCO3- — Ca^2+ + CO3^2- + H2O + CO2

Ca^2+ + CO3^2- —-CaCO3

317
Q

overall, weathering of continental silicate rocks

A

CO2 + CaSiO3 — SiO2 + CaCO3

318
Q

weathering of silicate rocks is

A

a CO2 sink

319
Q

silicate weathering followed by carbonate deposition causes a

A

drawdown of CO2

320
Q

weathering rates depend on

A

temperature
[H2CO3]_rain
supply of H2CO3

321
Q

temperature for weathering is a function of

A

pCO2

322
Q

[H2CO3] of rain (for weathering) is a function of

A

pCO2

323
Q

supply of H2CO3 for weathering is a function of

A

rainfall rate, which is a function of T

324
Q

** weathering rates directly depend on

A

pCO2

Temperature

325
Q

silicate weathering cycle feedback

A

negative feedback on T over geologic time
volcanic/meta. CO2 flux–> atmos.CO2–> surface T–>silicate weathering + carbonate deposition–. atmos CO2
also a line from atmos.CO2–>silicate weathering

326
Q

reverse of weathering

A

carbonate metamorphism

CaCO3 + SiO2 –> CaSiO3 + CO2

327
Q

in silicate weathering feedback what are the arrows from surface T to silicate weathering/carbonate deposition

A

reaction rate

rainfall rate

328
Q

in silicate weathering feedback what is arrow from atmospheric CO2–>silicate weathering/carbonate deposition

A

[H2CO3]_rain

329
Q

carbonate metamorphism depends on

A

amount of CaCO3 rock
tectonism (uplift)
volcanism

330
Q

carbonate metamorphism does not depend on

A

pCO2
T_surf
no feedback

331
Q

silicate weathering feedback is also called

A

WHAK feedback

332
Q

WHAK

A

Walker, Hayes, and Kasting, 1981

333
Q

silicate weathering timescales

A 
τ = AO carbon / silicate weathering
τ = 40,000 PgC / 0.03PgC/yr
τ = 4/3 x 10^6 yr
334
Q

based on timescales silicate weathering is important for

A

10^5 yr scales

1 000 000

335
Q

other factors affecting weathering rates

A 
rock type
changes in uplift rate
biotic enhancement of weathering
biotic control
amount of shelf seas
location of CaCO3 deposits
spreading/volcanism rates
336
Q

biological pump schematic

A

surface: DIC –via NPP—-C_org—via Respiration—DIC
C_org– Out of surface ocean
deep ocean: from C_org in surface— EP–burial flux (down)
EP— DIC— back up to surface ocean via upwelling
Burial flux—- C_org buried (down into sed.)
Burial flux—- back up to DIC (respiration)

337
Q

biological pump changes in surface ocean

A

CO2 + H2O – CH2O + O2
C_org fixed removes
1 Unit DIC
0 Unit Alk

338
Q

biological pump changes in deep ocean

A

adds
1 Unit DIC
0 Unit Alk

339
Q

biological pump compared to carbonate pump

A

like: more DIC moved to deep ocean
unlike: increases DIC relative to Alk (no change in Alk)

340
Q

biological pump moving DIC but not Alk

A

acidifies deep ocean
enhance CaCO3 dissolution
raises CCD

341
Q

CaCO3 dissolution is especially enhanced (by biological pump)

A

in top of sediment column (~few cm’s)

342
Q

organic carbon and oxygen

A

are intrinsically linked

added and removed by a closed biological cycle

343
Q

oxygen variability in different ocean layers

A

surface- source
deep- sink
sediment- sink

344
Q

surface oxygen source

A

oxygenic photosynthesis

345
Q

deep ocean oxygen sink

A

aerobic respiration

346
Q

ocean sediment oxygen sink

A

aerobic respiration (only if O2 available)

347
Q

spatial variability of oxygen forms

A

gradients

348
Q

low O2 supple or high C_org flux

A

anoxia (all O2 used)

349
Q

anoxia leads to

A

enhance C_org burial- anaerobic respiration is less effective

350
Q

anoxia example

A

saanich inlet- deep sill limits cycling of water, oxic layer overlays anoxic layer, bottom sediment is dark C_org

351
Q

*** major long term O2 source

A

burial of organic carbon

352
Q

burial or organic carbon

A

CO2 + H20 – CH2O + O2
CH2O buried- form rock
O2 left over— atmosphere

353
Q

O2 in C_org burial

A

accumulates as by-product

354
Q

** long term oxygen sink

A

oxidative weathering or organic sedimentary rock

355
Q

carbon fluxes that affect oxygen

A

GPP 100GtC/yr– 99.9% respired (big flux)

Burial 0.1GtC/yr– allows accumulation (small flux, big change over geologic time)

356
Q

what happens to sediment

A

deposited in ocean crust

deposited/accreted on continent

357
Q

carbon deposited on ocean crust

A

deposited on ocean crust— sea floor spreading, ocean crust to trench— subduct— carbon into mantle
trench-~- scrape off sediment– part of continents
subduct-~- devolatilize slab— c to atmos. as CO2/CO via arc volcano

358
Q

carbon deposited on continents

A

deposited/accreted on continent— tectonic uplift— weathering— carbon to atmos/ocean system
uplift-~- metamorphism– carbon to atmos./ocean system

359
Q

weathering of silicate rocks is a

A

CO2 sink

360
Q

CO2 drawdown

A

silicate weathering followed by carbonate deposition

361
Q

weathering rate depends on

A

temperature
[H2CO3]
supply of H2CO3

362
Q

temperature is a function of

A

pCO2

363
Q

[H2CO3]_rain is a function of

A

pCO2

364
Q

supply of H2CO3 is a function of

A

rainfall rate… function of temperature

365
Q

weathering rate directly depends on

A

pCO2 and temperature

366
Q

geology and climate

A

are inherently linked over geologic time

silicate weathering cycle is (-) feedback on T over geologic time

367
Q

reverse of weathering

A

carbonate metamorphism

368
Q

carbonate metamorphism

A

CaCO3 + SiO2 — CaSiO3 + CO2

369
Q

carbonate metamorphism depends on

A

amount of CaCO3 rock
tectonism (uplift)
volcanism

370
Q

carbonate metamorphism does not depend on

A

pCO2

T_surf

371
Q

silicate weathering feedback

A

volcanic/metamorphic CO2 flux–> atmos. CO2–> surface T–> silicate weathering followed by carbonate deposition–. atmos. CO2
also.. atmos CO2–> silicate weathering

372
Q

in silicate weathering feedback what is the surface T —> silicate weathering

A
  1. reaction rate

2. rainfall rate

373
Q

in silicate weathering feedback what is the atmospheric CO2 —> silicate weathering

A

[H2CO3]_rain

374
Q

silicate weathering feedback is called

A

WHAK feedback

375
Q

WHAK feedback

A

Walker, Hayes, and Kasting, 1981

376
Q

residence time

A

tau = AO carbon / silicate weathering = 40,000 PgC / 0.03 PgC/yr = 4/3 x10^6 yr

377
Q

other factors affecting weathering rate

A

rock type
changes in uplift rate
biotic enhancement of weathering

378
Q

how rock types affect weathering

A

Basalt weathers easily- unstable mineral structure at surface P,T
granite doesnt

379
Q

how changes in uplift rate affect weathering rate

A

supply of weatherable rock

380
Q

how biotic enhancement affects weathering rate

A

plants acidifying soils

381
Q

factors affection carbonate precipitation

A

biotic control

amount of shelf seas

382
Q

biotic control of carbonate precipitation

A

Ω: 3 vs. Ω: 30

biogenic precipitation or authigenic

383
Q

factors affecting CO2 flux

A

location of CaCO3 deposits

spreading/volcanisms rates

384
Q

cenozoic basics

A

65ma
good, detailed records
rise of mammals

385
Q

major events at bottom of cenozoic

A

K-T extinction (dinosaurs)

386
Q

boundaries in geologic time are usually set by

A

biostratigraphy

387
Q

earth 65ma

A

Europe fragmented
India in Indian ocean
Australia very south

388
Q

earth 50ma

A

atlantic opens more

India moves north

389
Q

earth 35ma

A

collisions between India/Asia begin

Australia moves north

390
Q

earth 20ma

A

drake passage opens
India ‘in place’
lower sea level
Europe closing up

391
Q

drake passage

A

area between SA and antartica

392
Q

earth now

A

Panama shut

very clear circumpolar seaway

393
Q

Eocene

A

last hothouse climate

start to get antarctic glaciation near end

394
Q

deconvolve ice volume

A

subtracting ice volumes to get T changes

relating deep ocean T to surface ocean T

395
Q

deep ocean temperature tells

A

mostly about polar temperatures

really cold bottom waters formed underneath ice shelf

396
Q

determining past climate from stomata

A

bigger stomata = lower CO2 levels

397
Q

warm epochs

A

paleocene
eocene
high levels CO2

398
Q

cool epochs

A

oligocene
miocene
pliocene
CO2 drop from Eocene - Oligocene

399
Q

~20ma cooling from

A

india ramming into asia– himalayans– monsoons and weathering

400
Q

strontium isotopes

A

87Sr / 86Sr

401
Q

87Sr

A

daughter product of rubidium, accumulates in continental rocks

402
Q

increase in 87Sr

A

increased weathering of continental rocks

403
Q

Tunguska treefall

A

Siberia, radial downfall of trees 40-50km, trees in middle didn’t fall

404
Q

why didn’t trees in middle of impact zone fall down

A

blast wave goes straight down under ‘bomb’, doesn’t push trees over
on edges wave pushes trees outwards

405
Q

why was there a blast wave

A

exploding rock increases T and PV, area surrounding rock is intensified in pressure, Heating causes rapid expansion; the air explodes outwards as blast wave

406
Q

ideal gas law

A

pV = nRT

407
Q

The velocity of a meteorite

A

from summing the escape velocity and some fraction of the Earth’s orbital velocity in quadrature

408
Q

velocity of meteorite, v =

A

√ v^2 + ( 1/a . v_orbit )^2

409
Q

v escape =

A

√ 2GM / r

M into page

410
Q

v orbit =

A

√ GM / d

M out of page

411
Q

typical a

A

2

412
Q

typical impact velocity

A

19 km/s

413
Q

specific energy of a meteor

A

e_k = 1/2 . v^2

414
Q

impacts compared to explosives

A

typical impact is 41X the specific energy of an explosive
impact = 1.7x10^8 J kg
TNT = 4.2x10^6 L kg

415
Q

simple crater

A

one impact zone with fracturing, breccia, impact melt, and impact ejecta on sides of crater

416
Q

complex crater

A

central peak uplift in middle

also fracturing, breccia, impact melt, impact ejecta on sides, sides are rougher

417
Q

why 1/a in velocity equation

A

what fraction of orbit speed is taken into account

ex. if rock is going to earth along earths orbit, orbit speed is negligible

418
Q

arizona crater

A

best preservation

419
Q

craters usually not well preserved

A

plate tectonics
vegetation
hydrologic cycle/weathering

420
Q

central peak

A

rebound from impact

421
Q

crater rim

A

excavated material

422
Q

knowing amount of material excavated in crater

A

allows estimate of energy of impact and mass of bolide

423
Q

Yucatan Peninsula, Mexico

A

Chicxulub crater (hit ground)

424
Q

using craters for dating

A

amount of craters vs. how often craters occur

425
Q

within 10 sec of impact

A

vaporization of impact material

local/regional blast damage

426
Q

within 1 min of impact

A

fires set by fireball

427
Q

within 10min of impact

A

ejecta, burial

earthquake, landslides, tsunami

428
Q

within 1 hour of impact

A

ejecta enter atmosphere

429
Q

hours-year of impact

A 
worldwide wildfires, smoke
total darkness
impact winter/ global cooling
ozone destruction by NOx
acid rain
geochemistry changes, nothing can grow
430
Q

years-millenia after impact

A

higher CO2- global warming
acid rain, warming– faster weathering
mass extinctions

431
Q

impact winter

A

massive amounts of dust in atmosphere, absorb all incoming sunlight

432
Q

NEOs

A

nearth-earth objects

433
Q

Power law

A

population of NEOs used to estimate frequencies of collision sizes

434
Q

power law graph

A

cumulative probability of impact/ yr (higher probability at top) vs. impact energy
decreasing graph
other y axis: average interval between impacts (yrs, decreasing up)

435
Q

what breaks down ozone

A

water- water doesn’t really make it to stratosphere because troposphere cools with altitude

436
Q

extinctions are usually from

A

consequences of impact, not impact itself

437
Q

large impact basin on moon

A

Orientale, Imbrium

438
Q

early bombardment era

A

more frequent, bigger impacts

lots of stray bodies floating around in early solar system

439
Q

late heavy bombardment

A

as evidenced by impact melts in lunar rocks dating 3.8-3.9Ga

may have been 0-4 ocean vaporizing impacts on earth

440
Q

power law today vs. hadean

A

todays line is lower, power of 10 lower intercept

441
Q

energy of impact required to vaporize ocean

A

10^34 - 10^36

442
Q

energy of KT impact

A

~10^31

443
Q

energy of imbrium impact

A

~10^33

444
Q

vaporizing the ocean

A

vaporize water and rock– release greenhouse gases and steam = steam atmosphere– LWR can’t escape
in time, everything rains out and reforms ocean

445
Q

time for ocean evaporation

A

few months

446
Q

time for formation of clouds and total drying of surface

A

1000 years

447
Q

time for cloud tops to cool, rain, and reform ocean

A

2000 years

448
Q

Tunguska event features

A

peewee, R >40m, local effects, romantic sunsets

449
Q

K-T impact summary

A

large, R >10km, half os species extinct, T changes, fires, darkness, chemical changes

450
Q

Moon forming event summary

A

super colossal, R >2000km, melt planet, drive off volatiles, wipe out life on planet

451
Q

the late veneer

A

delivery of volatiles to earth due to impacts

ocean came from comets

452
Q

mass flux of meteorite material to earth

A

10-100tons per day (mainly micrometeorites)

may add condensation nuclei

453
Q

cretaceous time

A

145-55mya

454
Q

cretaceous paleogeography

A

opening and closing of tethys sea

~105mya NA, SA, and Africa start to become familiar looking

455
Q

superchron

A

A superchron is a polarity interval lasting at least 10 million years. There are two well-established superchrons, the Cretaceous Normal and the Kiaman

456
Q

transgression

A

sea level rise

457
Q

regression

A

sea level drop

458
Q

sea level 150-65mya

A

apparently theres a transgression?

459
Q

why did the magnetic field not switch directions for much of the cretaceous

A

geomagnetic dynamo- sucking heat across core/mantel boundary, hard to reverse poles

460
Q

how often do geomagnetic reversals typically occur

A

in the last 20ma every 200,000- 300,000yr, although it has been more than twice that long since the last reversal
In the last 10 million years, there have been, on average, 4 or 5 reversals per million years

461
Q

superplume

A

occurs when a large mantle upwelling is convected to the Earth’s surface
should not be confused with a hot-spot- a superplume forms at the mantle-core boundary while a hot-spot occurs at the mantle-crust layer
create cataclysmic events, affect whole world when explode

462
Q

geodynamo and core mantle heat flux

A

continental scale superplume

huge areas of marine volcanism

463
Q

super plume needs a

A

superchron (Big Black Bar)

464
Q

why was there a transgression

A

young hot crust- buoyant and elevated = more mantle upwelling = lower continent
lower continent + displaced H2O = sea level rise = transgression

465
Q

LIP

A

large igneous province

466
Q

large igneous province

A

extremely large accumulation of igneous rocks, including liquid rock (intrusive) or volcanic rock formations (extrusive)
hot magma extrudes from inside Earth and flows out, source of LIPs is mantle plumes or plate tectonics

467
Q

cretaceous LIPs

A

~10 at the pacific MOR.. when atlantic was ~newly formed

468
Q

cretaceous events

A

superplume + dynamo
Big Black Bar
fast spreading rate + transgression
extensive, massive chalk deposits

469
Q

equator to pole range no

A

equator: 27ºC
pole: -21ºC

470
Q

equator to pole difference cretaceous

A

equator: 42ºC
Pole: 18ºC

471
Q

cretaceous CO2

A

~2000ppmv

higher volcanism = higher CO2

472
Q

polar amplification

A

any change in net radiation balance = larger T change near poles than the planetary average
poles increase faster than equator

473
Q

Strontium isotopes

A

84Sr: 0.56%
86Sr: 9.86%
87Sr: 7.00%
88Sr: 82.58%

474
Q

heat transport

A

movement of energy from equator to poles

475
Q

cretaceous heat transport

A

must have been higher since poles were warmer

mostly due to ocean transport

476
Q

primordial isotope

A

nuclides found on the Earth that have existed in their current form since before Earth was formed
ex. all 4 Sr isotopes

477
Q

radiogenic Sr isotope

A

87Rb—- 87Sr

t1/2 = 4.88x10^10yr

478
Q

Sr will substitute for

A

Ca

479
Q

Rb is

A

lithophile

incompatible in mantle, found dominantly in continents

480
Q

changes in 87Sr/86Sr

A

increasing: weathering of continents
decreasing: more basalt creation
87Sr decay on cont., 86Sr seafloor

481
Q

if 87Sr/86Sr decreasing

A

more basalt creation = more mantle flux = rise in sea level

482
Q

sea level rise in geologic history 90-100ma

A

associated with increase in 87Sr/86Sr- not what would be expected
coastal weathering from inland seas, increased mid latitude precipitation

483
Q

cretaceous chalk deposits

A

high Alk by increased continental weathering (silicate weathering)

484
Q

why increased continental weathering in cretaceous

A

superplume pushes up sea level- increase silicate weathering- higher CaCO3 deposit- anoxic bottom waters leave thick organic layers

485
Q

anoxic events from

A

rise in temperature

loss of O2 through thermohaline circulation due to respiration

486
Q

anoxic events cause

A

thick organic rich layers

487
Q

respiration

A

CH2O + O2 – CO2 + H2O

488
Q

carbon burial efficiency =

A

burial flux / depositional flux
using O2 buries C
high burial efficiency = low O2

489
Q

OAE

A

ocean anoxic events

geologically short, <1Myr

490
Q

contributors of OAE

A 
warm ocean, low O2 solubility
high C burial efficiency
high P supply from weathering/flooding
nutrient trapping in enclosed basins
Enhanced P regeneration under anoxic bottom waters
491
Q

Redfield ratio

A

C:N:P = 106:16:1

492
Q

sources of C,N,P

A

C,N have atmos. cycles, more available

P only source is cont. weathering

493
Q

why does increased P contribute to anoxia

A

it is generally the limiting resource ( + feedback)

494
Q

P burial efficiency

A

lower when anoxic

495
Q

EECO

A

early eocene climate optimum

496
Q

climate optimum

A

broad scale warm climates

high CO2

497
Q

after EECO

A

oligocene glaciation—sea level drop– MMCO

498
Q

basalt weathering

A

2-10X faster than granite- less stable minerals at surface conditions

499
Q

most important if basalt is

A

in equatorial/ITC region– hot and wet– lots of weathering

500
Q

ITC location

A

moves throughout year

501
Q

Deccan traps

A

basalt traps in India, 120Ma in SH, 65-30Ma migrate through ITC- major weathering

502
Q

movement of Deccan traps through ITC represents

A

Drift-weathering hypothesis

503
Q

why are Ethiopia traps less of a CO2 draw down

A

right on ITC but E side of continent– drier climate

504
Q

peak in basaltic + mixed blocks

A

~40-60Ma, also peak in all land mass, and granitic + sedimentary crust
consistent with EECO peek
hypothesis only considers CO2 sink (increased weathering)

505
Q

processes in interpreting paleoclimate

A

isotope records
look at graphs
build models
plate tectonics

506
Q

residence time =

A

steady state reservoir size / flux

507
Q

drift hypothesis

A

weatherable rocks drifting through high weathering zone

508
Q

sign of an OAE

A

a lot of chalk (Dover)

509
Q

weird think about OAE feedback

A

burying Corg is usually a nutrient loss (would be a - feedback) but in an AE, nutrients are easier to ‘strip out’— + feeback

510
Q

when we talk about OAE’s we mean what part

A

whole ocean except top ~100m’s

EOS (7 decks)

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