EOS 260 Part II Flashcards
steady state
dx/dt = 0
no change in position with time
stable or unstable
unstable steady state
unstable to a small perturbation- like on top of a hill
starting close to steady state system will diverge away
d2x/dt2>0
non-steady behaviour
transient
stable steady state
starting close to steady state system will converge to it
stable to a small perturbation
d2x/dt2<0
studying steady states
helps us understand the system
two stable steady states
must be separated by an unstable steady state
bistable system
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)
forcing
input to or control parameter affecting the system
will affect the state of the system
not affected itself by the state of the system
examples of forcing
solar forcing- sunlight earth receives
putting consecutively steeper ramps under a chair- force it to tip over
feedback
how the system will respond to a small perturbation or to a change in forcing
2 kinds
stabilizing feedback
negative feedback
diminishes effect of a perturbation, makes change smaller
pushes system back to stable steady state
Couplings
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
feedback loops
+1 to positive feeback
-1 to negative feedback
combined effect by multiplying them
most important gas absorbing insolation in atmosphere
H20 vapour
flux density units
W/m^2 = J / m^2 s
incoming solar radiation
341 W/m^2
destabilizing feedback
positive feedback
enhances the effect of a perturbation, makes change larger
pushes system on to next stable steady state after passing unstable steady state
outgoing longwave radiation
239 W/m^2
reflected solar radiation
102 W/m^2
total planetary albedo
outgoing/incoming
102/341 =~ 0.28 ~ 0.3
insolation absorbed by surface
161 W/m^2
average sunlight over whole earth
S/4 ~ 341 W/m^2
S
solar constant
S ~ 1368 W/m^2
relatively constant
where does insolation go
absorbed by atmosphere
absorbed by surface
reflected by surface
reflected by clouds/atmosphere
net energy absorbed by earth
0.9 W/m^2
absorbing ~1 W/m^2 more than it is emitting
where does surface radiation go
thermals evapotranspiration emitted by atmosphere back radiation- emitted back to surface out of atmosphere- through atm. window
rate of temperature change
proportional to energy imbalance ∆F
inversely proport. to system mass x specific heat capacity
dT/dt = ∆F/mc
∆F x area (of earth)
sunlight in the atmosphere
reflected/deflected by clouds, atmospheric molecules
atmospheric molecules
why the sky is blue
area of earth
1.5x10^14 m^2
why did our dt calculation for 2k in the ocean underestimate the time it would take
we found 730yrs
our model assumes ocean mixing, which takes ~1000yrs
what can be seen from calculating dt of certain temperature change in the atmosphere vs. the ocean
ocean takes much longer to heat
global warming is dependent on the ocean
how many seconds in a year
60x60x24x365.25 = 31557600s
2 layers of the atmosphere
troposphere
stratosphere
troposphere
lower layer, T profile set by large scale convection, T decreasing with altitude
stratosphere
upper layer, minimal motion, T set by radiative absorption and emission (O3), increases with altitude
between troposphere and stratosphere
tropopause
before ozone was formed
stratosphere would be constant T with altitude
black body
absorbs all electromagnetic radiation- appears black
emits radiation dependent on wavelength and temperature
Planck function
describes radiative emission of a black body
Stefan-Boltzmann law
integration of the Planck function
F_BB = σ T^4
F is flux, σ is constant, T is in Kelvin
mean earth surface T and F
T = 289 K F_BB = 396 W/m^2
Wein’s displacement law
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
some example λ_max
earth - 10µm (IR)
sun - ~500nm (visible)
largely different spectral regions
visible radiation absorption
not absorbed strongly by atmosphere
most radiation absorbed is at the surface
thermal radiation absorption
absorbed strongly by atmosphere
epsilon
emissivity band of thermal radiation absorption
Earth energy balance
sunlight absorbed
thermal radiation emitted
sunlight absorbed
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
absorbing area
π r_e^2
emitting area
4πr_e^2
thermal radiation emitted
earth emits thermal energy from a spherical surface
emitting at all points on earths surface
effective temperature of emission
T_eff
if earth had no atmosphere/greenhouse effect
effective temperature equation
S/4 (1-alpha) = σ T_eff^4
S is solar constant = 1368 W/m^2
alpha is albed = 0.3
attenuation
is the gradual loss in intensity of any kind of flux through a medium
attenuation of radiation in atmosphere
Reflectivity (R), Absorptivity (A), Transmissivity (T)
R + A + T = 1
reflectivity is also called albedo
solar radiation in the atmosphere, attenuation assumptions
R = A = 0
T = 1
‘all insolation goes to E’s surface’
thermal radiation in the atmosphere, attenuation assumptions
R = 0
A ~ 0.75
T ~ 0.25
all IR is absorbed and transmitted, no albedo of IR
Kirchoff’s Law
A_λ = ε_λ
absorptivity = emissivity
only valid where the λ is the same on both sides
grey body
F_grey = ε σ T^4
some fraction of the planck function, a function of every λ
Radiative equilibrium equations
surface
S/4 (1 -alpha) + ε σ Ta^4 = σ Ts^4
atmosphere
ε σ Ts^4 = 2 ε σ Ta^4
rearranging equilibrium equations gives you
Ts^4 = 2Ta^4 surface T is always bigger
a greenhouse gas is
a gas that absorbs thermal radiation
2 most important greenhouse gases
CO2, H20
strength of the greenhouse effect felt by a gas
depends on logarithm of abundance
recall that an increase in 1 is an increased factor of 10 with logarithm
radiative forcing
change in net flux at tropopause (down minus up)
surface temperature changes is proportional to
radiative forcing
Earths surface T without greenhouse effect
~255K
greenhouse effect requires
absorption and re-radiation of thermal energy in atmosphere by greenhouse gases
that atmosphere is colder than surface
strongest radiative forcing change
if you add CH4 you’ll see a stronger effect than CO2 because the base value is so low, log properties, larger change
Climate feedbacks
Planck (temperature) feedback (-)
Water vapour feedback (+)
Ice-Albedo feedback (+)
Planck feedback
solar forcing–> Temp—>outgoing thermal rad—-.temp
increase T = more energy emitted = cools down
water vapour feedback
amplifies T change
solar F–>T—>Atmos. H2O vapour—>greenhouse effect—. outgoing thermal radiation
could lead to runaway greenhouse if Planck feedback stopped working
saturation vapour pressure
e_s is proportional to exp T
equilibrium between liquid and gas
increase water vapour = increase T
e_s is pH2O
greenhouse forcing depend on
change of logarithm of gas abundance
strength of water vapour forcing depends
linearly on temperature change
Ice-albedo feedback
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
albedo examples
fresh snow 70-90 sea ice 50-75 desert 25-40 Forest/Grass 10-25 Ocean <10 -70 depending on incident ray
quick changing feedback
ice-albedo
there is snow, or there isnt
tau
thermal optical depth
strength of greenhouse effect
tau = tau_CO2 + tau_H2O
steady state energy balance on daisy world (no atmosphere)
σ T^4 = (1 - alpha)F_s
global albedo is a weighted sum of bare ground, black, white
solar flux on daisyworld
`increases linearly with time
daisyworld albedos
bareground: alpha_bare = 0.5
black: alpha_b = 0.25
white: alpha_w = 0.75
daisies are mesophile
growth rate beat depends on local T
max at 22.5ºC, no growth above 40ºC or below 5ºC
daisy growth
(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
daisy death
gamma A
daisyworld model
dA_w/dt = A_w ( x beta - gamma) dA_b/dt = A_b ( x beta - gamma)
statement of Gaia hypothesis
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
what does Gaia hypothesis mean
life has controlling influence on physical/chemical climate
life changes atmosphere, maintains habitability through time
theory led to Earth System science
why earth is special
temperature and pressure allow all three phases of water
triple point of phase diagram
all 3 phases exist in equilibrium
pressure on y axis, temperature on x
cryosphere
the frozen water part of the Earth system
southern hemisphere snow
snowy year round, not a lot of year round change
northern hemisphere snow
highly variable, huge variation w/ season, ice albedo changes are a function of the NH
antarctic sea ice cycle
not a lot of variation due to Antarctic circumpolar current
arctic sea ice minimum
september
forming a glacier requires
accumulation of multi-yr ice on land
precipitation (snow) in winter and failure of this to melt fully in summer
consequence of temperature change with altitude in regards to snow/ice
mountains receive most precipitation
snowmelt least likely at high altitude, and on poleward facing slopes
mountain glaciers form first
glacier
valleys, follow topography
ice field
big glacier, still constricted by topography but ‘drapes’ topography, flow direction directed by topography
ice sheet
unrestricted by topography, covers it and goes where it wants
ice shelf
floats at edge of ice sheet/glacier
net ice build up
accumulation
net ice loss
ablation
above glacier equilibrium line
snow > melt
accumulation zone
below glacier equilibrium line
melt > snow
ablation zone
mountain glaciers in the absence of melt
can grow to form ice sheets
ice sheet-altitude feedback
growth of ice sheet–raises altitude of ice surface– surface temperature is colder– melt inhibited
positive feedback
ice sheet flow
from centre of ice dome (high pt.) outward
thicker ice sheet
more likely to slide at base
movement of ice as a function of base T
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
warming ice sheet
speed up flow–loss of altitude–warmer T–more melt
positive feedback
glacial records
glacial valleys, moraines, striations, tillites, loess
ice rafted debris
piece breaks off of iceberg–flows away– starts melting– drops sediment– can deform bottom sediment
glacial loess
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
large ice cap instability
if ice reaches mid latitudes positive feedback will lead to global glaciation
from O-D energy balance model with ice-albedo feedback
small ice cap instability
if ice decreases too much it will all melt, no glaciation
ice sheets don’t get smaller, they collapse
north american ice sheets
lost: Laurentide, Cordillian, Scandanavian
remain: greenland
Phanerozoic/Proterozoic glaciations
Oligocene-Present (poles) Devonian, Cretaceous, Permian Ordovician- Siluria Cryogenian (Neoproterozoic) Siderian (paleoproterozoic)
faint young sun paradox
with lower S glaciation would be deep (with todays atmospheric composition)
yet geological evidence shows less glaciation earlier in history
stronger greenhouse?
isotope
same number of protons, different number of neutrons, same atomic number, different mass, same chemical properties, different physical properties
oxygen isotopes
16O - 99.76%
17O - 0.04%
18O - 0.2%
16,18 most commonly used for palaeoclimate
hydrogen isotopes
1H - 99.984%
2H - 0.016%
2H
deuterium ‘D’
isotopologues
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
most common water isotopologues
(1H)2(16O)
(1H)2(18O)
(1H)(2H)(16O) - HDO
in order of commonality
isotopic ratio
R = rare X / common X
ex. D: R = D/H = 0.016/99.984 = 0.00016
problem with isotope ratio
variations are very small numbers
delta notation ratio
δX = 1000 ((Rsample-Rstandard)/Rstandard)
units are ‘permil’ parts per thousand
δD =
1000 [ ( (D/H)sample - (D/H)smow ) / (D/H)smow ]
SMOW
standard mean ocean water
standard for H and O for ice sheets
isotopic mass balance
m_tot = m_1 + m_2
m_tot δ_tot = m_1 δ_1 + m_2 δ_2
δ18Osmow =
0%o
negative delta isotope values
depleted in 18O
preferentially left behind
the more negative, the colder
saturation vapour pressure
describes amount of water vapour in equilibrium with liquid water
saturation vapour pressure of isotopologues
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)
water vapour isotopes compared to ocean water
water vapour is isotopically light
δ18O_wv < δ18O_ocean
δD_wv < δD_ocean
precipitation is isotopically heavy relative to water vapour
δ18O_wv < δ18O_precip
δD_wv < δD_precip
as temperature decreases from the source region of moisture
increasingly higher fractions of the atmospheric water vapour have precipitated out
remaining water becomes increasingly isotopically light and further precipitation will be lighter
isotopic concentration of precipitate a function of
temperature at which precipitation occurs
precipitation is isotopically lighter
further from the source
at higher altitudes
why ice cores are drilled at the summit of an ice cap
reduces distortion due to ice flow
stratigraphy of ice core with depth
~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
ways to date stratigraphic layers
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
trade off between time resolution of ice core and length of record
higher snow accumulation rate = better time resolution
faster accumulation = faster flow, shorter time record
longer record, lower resolution ex.greenland
why layers are thinner down ice core
pressure
why might δD be better than δ18O in ice
mass differences
H 1/2 (half of mass)
O 16/18
δD versus age graph
less (-), warmer T’s, interglacials
glacials last longer than interglacials
packed snow
firn
bubble formation in glacier
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
age of air in ice
younger than ice itself (because they take so long to close)
and they close at different times
pacemaker of ice-ages
insolation changes (which affect ice volumes)
isotopes of ice sheets
accumulation of ice sheet removes light isotopes
ocean water become isotopically heavy
foraminifera
CaCO3 tests, must determine benthic vs. planktic
how forams represent ice volume
18O concentrated in CaCO3 relative to water (more with colder water), bust be planktonic (benthic waters don’t show much T variation)
marine δ18O graph
y-axis is opposite- heavy = high ice volume = cold
Pleistocen glaciation
regular glacial-interglacial cycles
before 700ka periodicity = 40ka
after 700ka periodicity = 100ka
change in glacial-interglacial periodicity
mid-Pleistocene climate transition
why are glacial studies better from antarctic
right on the pole-harder to melt ice sheet, more information contained in that
arctic ice is away from the pole, higher insolation
Croll-Milankovitch theory
Eccentricity
Axial tilt
Precession
eccentricity
how elliptical earths orbit it
eccentricity periodicity
100,000yrs
what eccentricity does
changes distribution of solar energy throughout year
perihelion
closest to sun (right now in NH Jan)
earth moving faster at this point of orbit to cover same amount of area
axial tilt
obliquity, the degree to which the axis is tilted changes
axial tilt periodicity
41,000yrs
axial tilt implications
changes strength of season
aphelion
earth farthest from sun
if obliquity were 0
no seasons
precession
wobble, change in where N/S axis points
precession periodicity
23,000 yrs
precession implications
relative strength and length of seasons
Croll-Milankovitch theory explains timing of glaciations
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
organic carbon dominantly fixed by
oxygenic photosynthesis
CO2 + H2O + hv —- CH2O + O2
inorganic matter reduced to organic matter
aerobic respiration
CH2O + O2 — CO2 + H2O
photosynthesis + aerobic respiration
closed cycle
decomposition of organic carbon in oxygen low environment
fermentation followed by methanogenesis
net: 2CH2O — CO2 + CH4
inorganic carbon
HCO3 -
CO3 2-
H2CO3
counter reaction to fermentation/methanogenesis
CH4 + 2O2 – CO2 + 2H2O
again, closed system
carbon reservoir cycling
biological reservoirs cycle quick
geochemical reservoirs cycle slow
atmospheric CO2 reservoir
760
photosynthesis takes 60 out
60 goes back in, fast cycling
residence time
tau = steady state reservoir size / flux
ex. 760 / 60 = ~13 yrs
basic structure of the ocean
2 layers: Mixed layer, deep ocean
mixed layer
~100m, geochemical equilibrium with atmosphere, well mixed by wind
ocean gradients
thermocline
chemocline
light penetration
thermocline
sudden small T decrease then, decreases gradually with depth down to ~100m where it reaches the minimum and stays ~constant
chemocline
ex. nutrient availability, abrupt decrease with depth
light penetration gradient
gradual decrease, area light penetrates = photic zone
Henry’s law
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)
1atm =
101325 pa
gases are more soluble in
cold water
DIC reservoirs, smallest to biggest
H2CO3
CO2
CO3 2-
HCO3 -
carbonate chemistry
collective chemistry of DIC species
conserved oceanic carbon quantities
DIC
Alkalinity
DIC =
[CO2] + [H2CO3] + [HCO3 -] + [CO3 2-]
Alkalinity
expresses charge balance, net + charge from conservative ions balanced by net - charge of weak acids
circumneutral conditions in ocean
pH > 4
pH =
- log [H+]
conservative ions
ions that don’t do much chemistry at ciarcumneutral conditions in ocean, salts, net +
[Na+] - [Cl-] + 2[Mg 2+] + 2[Ca 2+] > 0
Alkalinity, Alk =
[HCO3 -] + 2[CO3 2-] - [H+]
DIC reactions
CO2 + H20 ——– H2CO3
H2CO3 ——– H+ + HCO3 -
HCO3 - ——– H+ + CO3 2-
increasing pCO2
increases [CO2] – increases [H2CO3] — increases [H+] and [HCO3-] — increases [HCO3-] —- decreases [CO3 2-]
increased hydrogen ions, decreases pH, ocean acidification
why do we subract [Cl-] in conservative ion equation
because we are adding up the positive charges
pCO2 in DIC vs. Alkalinity graph
below ~where we are now (~0) - not possible
increasing in DIC = increased pCO2 (up to ~6µatm)
CO3 ^2- in DIC vs. Alkalinity graph
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
HCO3 ^- in DIC vs Alkalinity graph
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
CO2 in DIC vs. alkalinity graph
proportion of DIC that is CO2
currently 0
increases gradiently above present value, most of graph space is 1
pH in DIC vs. alkalinity graph
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
where we are now in terms of DIC and alkalinity
DIC ~ Alkalinity ~ 10^3
If DIC were»_space; alkalinity
pCO2 - high (6µatm)
pH - low (3)
~all Carbon is in CO2
if alkalinity»_space; DIC
can’t be
this is the ‘impossible’ part of the graph
if alkalinity and DIC both maxed
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
why - [H+] in Alk equation
adding up the negative ions
** H+ is generally negligible anyway
max alkalinity
all carbon in carbonate
Alk = [HCO3] + 2[CO3] - [H+]
the 2 coefficient is key, how to maximize equation
if Alkalinity = DIC
all carbon in HCO3
alkalinity high relative to DIC
carbon mostly in carbonate
very little carbon in CO2
decrease in alkalinity : DIC
transition to CO2 dominance
weathering rocks
adds to DIC and alkalinity
adding CO2 from atmosphere
only increases DIC
alkalinity units
we have been saying M (concentrations), but its actually supposed to be measured in equivelants/L
solid calcium carbonate precipitation from ocean
Ca^2+ + CO3^2- ——- CaCO3(s)
solubility product
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
saturation product
Ω = [Ca^2+].[CO3^2-] / k_sp
Ω < 1
subsaturated
CaCO3 dissolves
Thermodynamic Theory
Ω < 1 subsaturated
Ω = 1 saturated
Ω > 1 supersaturated
Ω > 3 biological precipitation
Ω > 1
supersaturated
CaCO3 precipitates
Ω > 30
authigenic precipitation (abiotic)
CaCO3 polymorphs
calcite
aragonite
aragonite properties
k_sp: 6.5x10^-7 mol^2/kg^2
Ω: 4
calcite properties
k_sp: 4.3x10^-7
Ω: 6
more soluble CaCO3 polymorph
aragonite- lower Ω, get to saturation point at lower concentration
decrease in ocean [CO3 ^2-]
decrease in saturation product (Ω)
problem for calcifying organisms
k_sp,cal/arag =
( [Ca2+].[CO3^2-] )_sat,cal
Ω_cal/arag =
[Ca2+][CO3^2-] / k_sp,cal/arag
variations in k_sp dominated by
depth (pressure)
as pressure increases k_sp
increases (solubility increases)
solubility increase with depth
2X / 4km depth
biologically precipitated CaCO3
sinks when organisms die– sedimented or dissolved in deep
CCD
carbonate compensation depth
rain rate of CaCO3 = CaCO3 dissolution
no sediment below
preservation of CaCO3 in sediment
~20-30%
depth that Ω = 1
saturation horizon, Lysocline
CaCO3 starts to dissolve
lysocline graph
Ω relatively high in surface
decreasing to saturation horizon
stays at 1 with increasing depth until reaches CCD then decreases further to 0
change in [CO3^2-] dominate
over changes in [Ca2+]
surface [CO3^2-] =~
250µmol/kg
Atlantic depth characteristics
deep ocean [CO3^2-] - 113µmol/kg
saturation horizon - 4.5km
Indian ocean depth characteristics
deep ocean [CO3^2-] - 83µmol/kg
saturation horizon - 3km
North Pacific ocean depth characteristics
deep ocean [CO3^2-] - 70µmol/kg
saturation horizon - <1km (not much CaCO3 sediment)
more deep ocean [CO3^2-]
deeper saturation horizon
Ω = Ca x CO3 / ksp — increase CO3 = increase Ω
deepest saturation horizon of the 3 oceans
Atlantic
changes in surface when CaCO3 precipitates
alkalinity decreases by 2 units
DIC decreases by 1 unit
changes in deep ocean when CaCO3 dissolves
Alkalinity increases by 2 units
DIC increases by 1 unit
carbonate pump
moves CaCO3 from surface to depth
why ocean carbon reservoir is so large
more C is stored in deep ocean than there would be if ocean was in equilibrium with atmosphere
Walker 1983
carbonate sedimentation throughout earths history constrains oceans chemistry and atmospheric composition
DIC is fixed to organic carbon by
photosynthesis in surface ocean
photosynthesis
CO2 + H2O == CH2O + O2
total organic carbon
GPP
GPP
gross primary productivity
total organic carbon after respiration by producers
NPP
NPP
net primary productivity
carbon sinking out of photic zone due to organism death
export production
in deep ocean most sinking carbon is
respired
carbon in deep ocean that is buried in sediment
burial flux
burial flux percentage of carbon in deep ocean
0.1%
organic carbon buried in sediment
massive deposits
massive deposits
black shale
up to 10% organic carbon
how much organic carbon is needed to turn sediment black
only a few %
pockets of organic carbon
diffuse deposits, kerogen
organic carbon fluxes
GPP - 100GtC/yr
NPP - 50 GtC/yr
export production - 10 GtC/yr
Burial flux - 0.1 GtC/yr
organic carbon that makes it out of surface ocean into deep ocean
~10%
carbon isotope ratios
12C - 0.9893
13C - 0.0107
14C - 10^ -17
stable carbon isotopes
12C
13C
molecules with different isotopic compositions
isotopologue
ex. 13CO2
δ (‰) =
[ R_sample / R_standard - 1 ] x10^3
R_sample or R_standard =
heavy isotope / light isotope
ie. 13C/12C
carbon standard
Pee Dee Belemnite (PeeDee formation)
R [13C/12C] PDB
0.011237
δ 13C= 0‰
R [13C/12C] CO2_atmospheric
0.01114486
lower than ocean
( - ) isotope values
depleted
permil
no units
≠ ppt (mass/volume)
δ 13C _CO2, atmospheric
-8.2‰
δ 13C_organic
~ -25‰
-20 - -40‰
δ 13C_carbonates
~ 0‰
δ 13C_HCO3-
~ 0‰
δ 13C_volcano
-5‰
δ 13C_fossil fuel
very low
graph of δ13C vs time is a decreasing trend, almost exact opposite of CO2 accumulations
fractionate
vibrational movements/energy differs btw. isotopes
light isotope bonds are weaker relative to heavy
light isotopes are easier to evaporate relative to heavy
vibrational movement graph
# of atoms vs. velocity normal distribution, most atoms in mean area, some in tails tails of distribution are longer in light isotopes
fractionation steps in plants
diffusion
carboxylation
Diffusion
ε = 4.4‰
lighter molecule enters cell preferentially
cell depleted in 13C
carboxylation
ε = (-15) - (-30)‰
energetically expensive process
easier to break bond in light isotopes
ε
difference in δ values between 2 species
after 2 steps of fractionation
cell becomes ~ -25‰
but this depends on whether we are talking about land or ocean plants
why does overall plant cell depletion depend on land/ocean
source of original carbon that is taken into cell in step 1
air: -8.2‰
HCO3- : 0‰
surface carbon cycle
F_in = F_org + F_carb = F_out
F_in
volcanoes
-5‰
F_org
ΔB
-25‰
F_carb
δ 13HCO^3-
0‰
ΔB
biologic fractionation
f_org =
F_org / F_in
= δin - δcarb / δorg - δcarb
= ( -5 - 0 ) / ( -25 - 0 ‰ ) = 1/5 = 0.2
which F is stable through earths history
Fin
% of carbon buried as organic carbon
20%- preserves oxygenated state of atmosphere
increase in δ 13C_carb
more f_org
higher productivity
increase in f_org
burying alot of light carbon so δ13C increases because f_carb decreases, have to maintain isotopic balance
represent ordinary continental silicate rocks as
CaSiO3
ordinary continental silicate rocks
granite, basalt
carbonic acid in rain
2[CO2 + H2O — H2CO3]
continental rocks weathered by carbonic acid in rain
CaSiO3 + 2H2CO3 — Ca^2+ + 2HCO^3- + SiO2 + H2O
consequence of weathering continental rocks
draws down 2CO2 from atmosphere, transfers Ca,2HCO3 to ocean
continental rock weathering drawing down CO2
2CO2 removed from atmosphere = 2 units of DIC
continental rock weathering transferred ions
add 2 units of DIC
add 2 units of Alk
net changes from weathering of continental rock
ocean gain:
0 units DIC
2 units Alk
atmos. CO2 cancelled out by ocean CO2 (Henrys Law)
why can we call the atmospheric CO2 ~ ocean CO2, and cancel out gains and losses?
because they equilibrate so rapidly
continental rock weathering shifts carbon
from CO2 – CO3 ^ 2-
shifting carbon from CO2 to CO3 ^2- results in
increase in CaCO3 saturation
higher CaCO3 deposition
weathering continental rock, shifting carbon equations
Ca^2+ + 2HCO3- — Ca^2+ + CO3^2- + H2O + CO2
Ca^2+ + CO3^2- —-CaCO3
overall, weathering of continental silicate rocks
CO2 + CaSiO3 — SiO2 + CaCO3
weathering of silicate rocks is
a CO2 sink
silicate weathering followed by carbonate deposition causes a
drawdown of CO2
weathering rates depend on
temperature
[H2CO3]_rain
supply of H2CO3
temperature for weathering is a function of
pCO2
[H2CO3] of rain (for weathering) is a function of
pCO2
supply of H2CO3 for weathering is a function of
rainfall rate, which is a function of T
** weathering rates directly depend on
pCO2
Temperature
silicate weathering cycle feedback
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
reverse of weathering
carbonate metamorphism
CaCO3 + SiO2 –> CaSiO3 + CO2
in silicate weathering feedback what are the arrows from surface T to silicate weathering/carbonate deposition
reaction rate
rainfall rate
in silicate weathering feedback what is arrow from atmospheric CO2–>silicate weathering/carbonate deposition
[H2CO3]_rain
carbonate metamorphism depends on
amount of CaCO3 rock
tectonism (uplift)
volcanism
carbonate metamorphism does not depend on
pCO2
T_surf
no feedback
silicate weathering feedback is also called
WHAK feedback
WHAK
Walker, Hayes, and Kasting, 1981
silicate weathering timescales
τ = AO carbon / silicate weathering τ = 40,000 PgC / 0.03PgC/yr τ = 4/3 x 10^6 yr
based on timescales silicate weathering is important for
10^5 yr scales
1 000 000
other factors affecting weathering rates
rock type changes in uplift rate biotic enhancement of weathering biotic control amount of shelf seas location of CaCO3 deposits spreading/volcanism rates
biological pump schematic
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)
biological pump changes in surface ocean
CO2 + H2O – CH2O + O2
C_org fixed removes
1 Unit DIC
0 Unit Alk
biological pump changes in deep ocean
adds
1 Unit DIC
0 Unit Alk
biological pump compared to carbonate pump
like: more DIC moved to deep ocean
unlike: increases DIC relative to Alk (no change in Alk)
biological pump moving DIC but not Alk
acidifies deep ocean
enhance CaCO3 dissolution
raises CCD
CaCO3 dissolution is especially enhanced (by biological pump)
in top of sediment column (~few cm’s)
organic carbon and oxygen
are intrinsically linked
added and removed by a closed biological cycle
oxygen variability in different ocean layers
surface- source
deep- sink
sediment- sink
surface oxygen source
oxygenic photosynthesis
deep ocean oxygen sink
aerobic respiration
ocean sediment oxygen sink
aerobic respiration (only if O2 available)
spatial variability of oxygen forms
gradients
low O2 supple or high C_org flux
anoxia (all O2 used)
anoxia leads to
enhance C_org burial- anaerobic respiration is less effective
anoxia example
saanich inlet- deep sill limits cycling of water, oxic layer overlays anoxic layer, bottom sediment is dark C_org
*** major long term O2 source
burial of organic carbon
burial or organic carbon
CO2 + H20 – CH2O + O2
CH2O buried- form rock
O2 left over— atmosphere
O2 in C_org burial
accumulates as by-product
** long term oxygen sink
oxidative weathering or organic sedimentary rock
carbon fluxes that affect oxygen
GPP 100GtC/yr– 99.9% respired (big flux)
Burial 0.1GtC/yr– allows accumulation (small flux, big change over geologic time)
what happens to sediment
deposited in ocean crust
deposited/accreted on continent
carbon deposited on ocean crust
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
carbon deposited on continents
deposited/accreted on continent— tectonic uplift— weathering— carbon to atmos/ocean system
uplift-~- metamorphism– carbon to atmos./ocean system
weathering of silicate rocks is a
CO2 sink
CO2 drawdown
silicate weathering followed by carbonate deposition
weathering rate depends on
temperature
[H2CO3]
supply of H2CO3
temperature is a function of
pCO2
[H2CO3]_rain is a function of
pCO2
supply of H2CO3 is a function of
rainfall rate… function of temperature
weathering rate directly depends on
pCO2 and temperature
geology and climate
are inherently linked over geologic time
silicate weathering cycle is (-) feedback on T over geologic time
reverse of weathering
carbonate metamorphism
carbonate metamorphism
CaCO3 + SiO2 — CaSiO3 + CO2
carbonate metamorphism depends on
amount of CaCO3 rock
tectonism (uplift)
volcanism
carbonate metamorphism does not depend on
pCO2
T_surf
silicate weathering feedback
volcanic/metamorphic CO2 flux–> atmos. CO2–> surface T–> silicate weathering followed by carbonate deposition–. atmos. CO2
also.. atmos CO2–> silicate weathering
in silicate weathering feedback what is the surface T —> silicate weathering
- reaction rate
2. rainfall rate
in silicate weathering feedback what is the atmospheric CO2 —> silicate weathering
[H2CO3]_rain
silicate weathering feedback is called
WHAK feedback
WHAK feedback
Walker, Hayes, and Kasting, 1981
residence time
tau = AO carbon / silicate weathering = 40,000 PgC / 0.03 PgC/yr = 4/3 x10^6 yr
other factors affecting weathering rate
rock type
changes in uplift rate
biotic enhancement of weathering
how rock types affect weathering
Basalt weathers easily- unstable mineral structure at surface P,T
granite doesnt
how changes in uplift rate affect weathering rate
supply of weatherable rock
how biotic enhancement affects weathering rate
plants acidifying soils
factors affection carbonate precipitation
biotic control
amount of shelf seas
biotic control of carbonate precipitation
Ω: 3 vs. Ω: 30
biogenic precipitation or authigenic
factors affecting CO2 flux
location of CaCO3 deposits
spreading/volcanisms rates
cenozoic basics
65ma
good, detailed records
rise of mammals
major events at bottom of cenozoic
K-T extinction (dinosaurs)
boundaries in geologic time are usually set by
biostratigraphy
earth 65ma
Europe fragmented
India in Indian ocean
Australia very south
earth 50ma
atlantic opens more
India moves north
earth 35ma
collisions between India/Asia begin
Australia moves north
earth 20ma
drake passage opens
India ‘in place’
lower sea level
Europe closing up
drake passage
area between SA and antartica
earth now
Panama shut
very clear circumpolar seaway
Eocene
last hothouse climate
start to get antarctic glaciation near end
deconvolve ice volume
subtracting ice volumes to get T changes
relating deep ocean T to surface ocean T
deep ocean temperature tells
mostly about polar temperatures
really cold bottom waters formed underneath ice shelf
determining past climate from stomata
bigger stomata = lower CO2 levels
warm epochs
paleocene
eocene
high levels CO2
cool epochs
oligocene
miocene
pliocene
CO2 drop from Eocene - Oligocene
~20ma cooling from
india ramming into asia– himalayans– monsoons and weathering
strontium isotopes
87Sr / 86Sr
87Sr
daughter product of rubidium, accumulates in continental rocks
increase in 87Sr
increased weathering of continental rocks
Tunguska treefall
Siberia, radial downfall of trees 40-50km, trees in middle didn’t fall
why didn’t trees in middle of impact zone fall down
blast wave goes straight down under ‘bomb’, doesn’t push trees over
on edges wave pushes trees outwards
why was there a blast wave
exploding rock increases T and PV, area surrounding rock is intensified in pressure, Heating causes rapid expansion; the air explodes outwards as blast wave
ideal gas law
pV = nRT
The velocity of a meteorite
from summing the escape velocity and some fraction of the Earth’s orbital velocity in quadrature
velocity of meteorite, v =
√ v^2 + ( 1/a . v_orbit )^2
v escape =
√ 2GM / r
M into page
v orbit =
√ GM / d
M out of page
typical a
2
typical impact velocity
19 km/s
specific energy of a meteor
e_k = 1/2 . v^2
impacts compared to explosives
typical impact is 41X the specific energy of an explosive
impact = 1.7x10^8 J kg
TNT = 4.2x10^6 L kg
simple crater
one impact zone with fracturing, breccia, impact melt, and impact ejecta on sides of crater
complex crater
central peak uplift in middle
also fracturing, breccia, impact melt, impact ejecta on sides, sides are rougher
why 1/a in velocity equation
what fraction of orbit speed is taken into account
ex. if rock is going to earth along earths orbit, orbit speed is negligible
arizona crater
best preservation
craters usually not well preserved
plate tectonics
vegetation
hydrologic cycle/weathering
central peak
rebound from impact
crater rim
excavated material
knowing amount of material excavated in crater
allows estimate of energy of impact and mass of bolide
Yucatan Peninsula, Mexico
Chicxulub crater (hit ground)
using craters for dating
amount of craters vs. how often craters occur
within 10 sec of impact
vaporization of impact material
local/regional blast damage
within 1 min of impact
fires set by fireball
within 10min of impact
ejecta, burial
earthquake, landslides, tsunami
within 1 hour of impact
ejecta enter atmosphere
hours-year of impact
worldwide wildfires, smoke total darkness impact winter/ global cooling ozone destruction by NOx acid rain geochemistry changes, nothing can grow
years-millenia after impact
higher CO2- global warming
acid rain, warming– faster weathering
mass extinctions
impact winter
massive amounts of dust in atmosphere, absorb all incoming sunlight
NEOs
nearth-earth objects
Power law
population of NEOs used to estimate frequencies of collision sizes
power law graph
cumulative probability of impact/ yr (higher probability at top) vs. impact energy
decreasing graph
other y axis: average interval between impacts (yrs, decreasing up)
what breaks down ozone
water- water doesn’t really make it to stratosphere because troposphere cools with altitude
extinctions are usually from
consequences of impact, not impact itself
large impact basin on moon
Orientale, Imbrium
early bombardment era
more frequent, bigger impacts
lots of stray bodies floating around in early solar system
late heavy bombardment
as evidenced by impact melts in lunar rocks dating 3.8-3.9Ga
may have been 0-4 ocean vaporizing impacts on earth
power law today vs. hadean
todays line is lower, power of 10 lower intercept
energy of impact required to vaporize ocean
10^34 - 10^36
energy of KT impact
~10^31
energy of imbrium impact
~10^33
vaporizing the ocean
vaporize water and rock– release greenhouse gases and steam = steam atmosphere– LWR can’t escape
in time, everything rains out and reforms ocean
time for ocean evaporation
few months
time for formation of clouds and total drying of surface
1000 years
time for cloud tops to cool, rain, and reform ocean
2000 years
Tunguska event features
peewee, R >40m, local effects, romantic sunsets
K-T impact summary
large, R >10km, half os species extinct, T changes, fires, darkness, chemical changes
Moon forming event summary
super colossal, R >2000km, melt planet, drive off volatiles, wipe out life on planet
the late veneer
delivery of volatiles to earth due to impacts
ocean came from comets
mass flux of meteorite material to earth
10-100tons per day (mainly micrometeorites)
may add condensation nuclei
cretaceous time
145-55mya
cretaceous paleogeography
opening and closing of tethys sea
~105mya NA, SA, and Africa start to become familiar looking
superchron
A superchron is a polarity interval lasting at least 10 million years. There are two well-established superchrons, the Cretaceous Normal and the Kiaman
transgression
sea level rise
regression
sea level drop
sea level 150-65mya
apparently theres a transgression?
why did the magnetic field not switch directions for much of the cretaceous
geomagnetic dynamo- sucking heat across core/mantel boundary, hard to reverse poles
how often do geomagnetic reversals typically occur
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
superplume
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
geodynamo and core mantle heat flux
continental scale superplume
huge areas of marine volcanism
super plume needs a
superchron (Big Black Bar)
why was there a transgression
young hot crust- buoyant and elevated = more mantle upwelling = lower continent
lower continent + displaced H2O = sea level rise = transgression
LIP
large igneous province
large igneous province
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
cretaceous LIPs
~10 at the pacific MOR.. when atlantic was ~newly formed
cretaceous events
superplume + dynamo
Big Black Bar
fast spreading rate + transgression
extensive, massive chalk deposits
equator to pole range no
equator: 27ºC
pole: -21ºC
equator to pole difference cretaceous
equator: 42ºC
Pole: 18ºC
cretaceous CO2
~2000ppmv
higher volcanism = higher CO2
polar amplification
any change in net radiation balance = larger T change near poles than the planetary average
poles increase faster than equator
Strontium isotopes
84Sr: 0.56%
86Sr: 9.86%
87Sr: 7.00%
88Sr: 82.58%
heat transport
movement of energy from equator to poles
cretaceous heat transport
must have been higher since poles were warmer
mostly due to ocean transport
primordial isotope
nuclides found on the Earth that have existed in their current form since before Earth was formed
ex. all 4 Sr isotopes
radiogenic Sr isotope
87Rb—- 87Sr
t1/2 = 4.88x10^10yr
Sr will substitute for
Ca
Rb is
lithophile
incompatible in mantle, found dominantly in continents
changes in 87Sr/86Sr
increasing: weathering of continents
decreasing: more basalt creation
87Sr decay on cont., 86Sr seafloor
if 87Sr/86Sr decreasing
more basalt creation = more mantle flux = rise in sea level
sea level rise in geologic history 90-100ma
associated with increase in 87Sr/86Sr- not what would be expected
coastal weathering from inland seas, increased mid latitude precipitation
cretaceous chalk deposits
high Alk by increased continental weathering (silicate weathering)
why increased continental weathering in cretaceous
superplume pushes up sea level- increase silicate weathering- higher CaCO3 deposit- anoxic bottom waters leave thick organic layers
anoxic events from
rise in temperature
loss of O2 through thermohaline circulation due to respiration
anoxic events cause
thick organic rich layers
respiration
CH2O + O2 – CO2 + H2O
carbon burial efficiency =
burial flux / depositional flux
using O2 buries C
high burial efficiency = low O2
OAE
ocean anoxic events
geologically short, <1Myr
contributors of OAE
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
Redfield ratio
C:N:P = 106:16:1
sources of C,N,P
C,N have atmos. cycles, more available
P only source is cont. weathering
why does increased P contribute to anoxia
it is generally the limiting resource ( + feedback)
P burial efficiency
lower when anoxic
EECO
early eocene climate optimum
climate optimum
broad scale warm climates
high CO2
after EECO
oligocene glaciation—sea level drop– MMCO
basalt weathering
2-10X faster than granite- less stable minerals at surface conditions
most important if basalt is
in equatorial/ITC region– hot and wet– lots of weathering
ITC location
moves throughout year
Deccan traps
basalt traps in India, 120Ma in SH, 65-30Ma migrate through ITC- major weathering
movement of Deccan traps through ITC represents
Drift-weathering hypothesis
why are Ethiopia traps less of a CO2 draw down
right on ITC but E side of continent– drier climate
peak in basaltic + mixed blocks
~40-60Ma, also peak in all land mass, and granitic + sedimentary crust
consistent with EECO peek
hypothesis only considers CO2 sink (increased weathering)
processes in interpreting paleoclimate
isotope records
look at graphs
build models
plate tectonics
residence time =
steady state reservoir size / flux
drift hypothesis
weatherable rocks drifting through high weathering zone
sign of an OAE
a lot of chalk (Dover)
weird think about OAE feedback
burying Corg is usually a nutrient loss (would be a - feedback) but in an AE, nutrients are easier to ‘strip out’— + feeback
when we talk about OAE’s we mean what part
whole ocean except top ~100m’s
