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Earth System Science II - The Oceans as a Chemical Reservoir.
The relative abundance of the elements composing the Earth is largely similar to that of the Sun and the Solar Nebula, and the Earth's `major elements' are therefore in order of decreasing valence or ionic charge (with the exception of phosphorus, sulphur, and carbon (P, S, C)):
Si, Ti,----Al, Fe3+,----Fe2+, Mn, Mg, Ca,----Na, K----P, S, CElemental Cosmic Abundances; atoms per 10^6 Si.
As a differentiated body, most
of the Earth's iron and nickel (plus some sulphur and/or oxygen)
is contained in the Earth's core, whereas the uppermost 100 km of the
Earth's mantle is largely composed of oxygen, magnesium, and silicon in
the approximate ratio 3.5:1.5:1, corresponding to a mixture of the
minerals olivine (58%), orthopyroxene (30%), clinopyroxene (10%),
and spinel (2%). Most of the major elements other than Si and Mg
are contained in the clinopyroxene and spinel minerals, and the removal
of these elements from the mantle to the crust and eventually to the
oceans is initiated by the melting of the pyroxene and spinel to form basaltic
magma during the decompression melting
events described in the previous lecture. The melting removes virtually
all of the Al, Ca, Na, K, P, and CO2 and other gases from the mantle
material.
The basaltic magma is
further enriched in the major elements by the gravitational separation
and removal of magnesium minerals from the melt as it cools to form
oceanic crust at the mid-ocean ridge, and during this stage the basalt
is particularly enriched in Si to the point that it is capable of
crystallizing in the form of quartz. Gases associated with the volcanic
materal are also released at this time to the overlying ocean and
atmosphere, and the chemical composition of parts of the basaltic oceanic
crust is modified at the mid-ocean ridge stage by reaction with sea-water
introduced by the `black and white smoker' hydrous convection cells.
The modification particularly involves the introduction of oxygen
into the system, thereby fixing iron in its trivalent state, the
hydration of the basaltic minerals, the addition of potassium, the
remobilization of magnesium and sulphur, and the removal of part of any
residual halogens (chlorine) and carbon dioxide.
During the subduction
stage of the cycle, the dehydration of the hydrated oceanic crust
releases water into the overlying mantle. At the high pressure and
temperature involved at this stage, the fluid is able to dissolve large
concentrations of the major elements, including even the relatively
immobile element Ti, which are therefore enriched in the mantle melt
that eventually rises to feed the overlying island arc volcanic system.
Since the arc volcanics are relatively enriched in Si and K, fractionation
of the arc melts can lead to the generation of typically continental K-
feldspar bearing granitoids and effusive rhyolites. Such rocks are never
found associated with oceanic crust. If the arcs are built on pre-existing
continental crust such as the Andes, the arc melts may become further
enriched by the incorporation of the light element component (Na,
K) contained in the crust. In this way the upper crust is selectively
enriched
in the alkali feldspar minerals and quartz typically found in granites.
Further upwards displacement
and concentration of Si and the alkali and alkali earth elements,
and downward relative displacement of magnesium and iron, will also
take place during remelting of arc and preexisting continental rocks during
continental collision.
The overall effect of
the mantle to arc/collision transfer of material is therefore the
development of continental crust enriched in Si, Al, Ca, Na, K, and P.
The Mg, Fe, Ti, and Mn are recycled back into the mantle, and Cl,
CO2 and other gases are added directly to the hydrosphere and atmosphere
during all stages of the mantle to continent cycle, with a particular
boost being given to CO2 during collisional events due to decarbonization
of those sedimentary carbonate rocks forced deep into the crust. Some of
the processes involved in the large scale transfer of elements between
the the mantle, crust, oceans, and atmosphere are illustrated in the following
diagram:
The Mantle - Crust - Ocean - Mantle cycle - the Earth’s geochemical and biological reservoirs.
The transfer processes
are complicated, and in the case of Carbon, Berner (1999) recognises both
short-term residence and long-term residence cycles. The short-term cycle
involves 1) gas exchange between the atmosphere
and the oceans, 2) accumulation of biotic
carbon in soils, and subsequent transfer of
the carbon to the atmosphere and oceans by degassing and river transport,
respectively, and 3) direct carbon exchange between the biota
and the atmosphere by terrestrial photosynthesis and respiration.
The long-term cycle involves the various short-term
cycle reservoirs and the long term carbonate
and sedimentary carbon reservoirs. In
this case carbon from the short term reservoirs is transfered to the long
term carbonate and sedimentary carbon reservoirs by burial,
and returned from the long term reservoirs to the short term reservoirs
by weathering, and volcanic
(mantle source), diagenetic, and subduction/collisional
degassing.
The
rate at which carbon is transferred between the various reservoirs, and
the bulk carbon content of the reservoirs is subject to the prevailing
physical conditions, and the recycling behaviour of other elements. For
example, carbon in the biotic reservoir is in part controlled by the availability
of nutrients, in particular phosphorus and iron, which is in turn controlled
by ocean currents which are themselves controlled in a complex manner by
the climate. An increase in oxygen would cause a decrease in available
phosphorus due to burial of ferric oxides and adsorbed phosphate. This
would in turn sequentially cause a decrease in the biota, and a decrease
in the amount of oxygen formed by photosynthesis. This is termed a negative
feedback. On the other hand an increase
in oxygen might cause fewer land plants (forest fires) and therefore less
weathering of organic matter, leading in turn to a further relative increase
in oxygen. This is termed a positive feedback.
Recommended reading:
Berner, R., 1999, A new look at the long-term carbon cycle, GSA-Today,
v. 9, no. 11, p. 1-6.
CO2 AS A WEATHERING AGENT
The major elements brought
to the surface of the continents from the mantle are subsequently transferred
to the oceans thought the simple mechanism of rain water weathering
and riverine transport of the dissolved rock components.
The main agent in the
weathering process is CO2 dissolved in rain water to form carbonic acid,
and the essential (non-equilibrium) reaction relationships are:
albite, orthoclase + water + carbon dioxide = kaolinite + alkali bicarbonate + silicic acid 2 (Na,K)AlSi3O8 + 11 H2O + 2CO2 = Al2Si2O5(OH)4 + 2(Na,K)+ 2HCO3- + 4 H4SiO4 and anorthite + water + carbon dioxide = kaolinite calcium bicarbonate CaAl2Si2O8 + 3 H2O + 2 CO2 = Al2Si2O5(OH)4 + Ca++ + 2HCO3--Weathering is the main feed back mechanism that keeps the level of atmospheric CO2 - which is continuously being removed from the mantle and added to the hydro-atmosphere in proportion to the amount of oceanic crust being formed at mid-ocean ridges - within the limits that allow the continuation of biological activity. For example, if global CO2 levels should rise, the Earth's atmosphere would warm, which would in turn increase precipitation and the rate of weathering, and also increase the amount of continental crust exposed to weathering as a result of the melting of the polar ice-caps. (This would be mitigated or even negated by the decrease in exposed continental crust as a result of flooding caused by the raise in sea level, which would itself be negated by the rise of the continents, e.g. North America, due to the removal of the weight of the ovelying ice.) The rate of consumption of CO2 by weathering may thus increase and the Earth would cool once again. Interestingly, if all the continents were to move to polar locations and be covered with ice, the average temperature of the Earth would increase because weathering as a control on CO2 levels would be insignificant. It has been estimated that halving the amount of emergent land area will elevate CO2 levels enough to raise land surface temperatures 10 deg.C, and vice versa, an increase in temperature sufficient to double the silicate weathering rate to keep long-term CO2 mantle degassing and silicate weathering balanced. This relationship between CO2 and the CaO released by weathering (the 2-10 formula) can be perturbed by changes in solar luminosity, orbital parameters, carbonate/carbon burial rates, mantle degassing rates, continent hypsometry, supercontinent formation, continent location, and orogenesis.
The CO2 - weathering feedback mechanism.
Recent studies (Vitouseck, P.M. et al. 1997. Soil and ecosystem development across the Hawaiian Islands. GSA Today, 7, 9, 1-8) of weathering effects on the volcanic islands of Hawaii have shown that it take only 20,000 years to reduce the alkali elements and Si to less than 10% of their initial quantities, and that by 150,000 years the original rock mass is decreased by 50%.
Weathering of Hawaiian volcanic rocks
THE BEHAVIOUR OF REE DURING WEATHERINGDuring weathering organic acid-charged rain water is capable of dissolving and slightly fractionating REE towards HREE enrichment in the residual clay. As the solutions percolate to lower horizons they lose their acidity and the REE are incorporated into newly formed clays, allowing enrichments up to a factor of 2. The overall concentration of REE in the soil profile is however not changed, and the REE in the sediments derived from the weathered rock remains essentially that of the unweathered rock, with most of the REE's concentrated in the clay fraction. In cases of extreme weathering (e.g. Amazon deep-sea muds), particle size sorting may however ind uce a slight LREE enrichment. It should be noted that where sandstones have accumulated heavy minerals with large abundances of REE such as zircon, monazite or allanite, the REE pattern of the sandstones may largely reflect the mixing ratios of these minerals.
% contribution of various
sources to Na content of river water (Berner and Berner,
1996)
cyclic salt silicates
evaporites pollution
8%
22%
42%
28%
Concentration of alkali
metals and Cl in River Water (Pinet, 1996)
ppm kg/kg
%
Ca2+
15 1.50 x10^-5 12.4
Na+
6.3 6.30 x10^-6 5.2
Na+ (22% weathering) 1.6
1.60 x10^-6
K+
2.3 2.30 x10^-6 1.9
Cl
7.8 7.80 x10^-6 6.5
Bicarbonate
58.8 5.88 x10^-5 48.7%
Concentration of alkali
metals and Cl in River Water and Seawater (Berner and Berner,
1996)
(Mass concentration = Atomic mass x micromoles/litre x 10^-9 kg/kg;
e.g. 558000 x 35.45 x 10^-9 = 0.0198 kg/kg = 1.98 x 10^-2 kg/kg)
Component River Water Seawater Tau(r)
microM/L kg/kg microM/L
kg/kg (1000
yr)
Ca++
367 1.5 x 10^-5 10,500
4.2 x10^-4 1,000
Na+
315 7.2 x 10^-6 479,000
1.1 x 10^-2 55,000
Na+ (weathering) 69 1.6 x 10^-6
K+
36 1.4 x 10^-6 10,400
4.07 x 10^-4 10,000
Cl-
230 8.1 x 10^-6 558,000
1.98 x 10^-2 87,000
The amounts of the major elements contained in the various crustal reservoirs are very large and and it is consequently often difficult to establish a mental image of the quantities involved. As an example, one might consider the question: How much salt is there in the oceans?
The surface area of the Earth = 4piR^2 = 4 x 3.1416
x (6378)^2 km = 5.1119 x 10^8
Continental area = 4piR^2 x 0.29 = 4 x
3.1416 x (6378)^2 km x 0.29 = 1.4824 x 10^8
Ocean area = 4piR^2 x 0.71 = 4 x
3.1416 x (6378)^2 km x 0.71 = 3.6294 x 10^8
Ratio of oceanic area to continental area = 2.4483
1 m = 100 cm;
1 km = 1000 m;
1 l = 1000 cm^3 = 1 dm^3; 1
km^3 = 109 m^3 = 10^12 dm^3 (litres) = 10^12 kg
(SG of water = 1; Density of water is 1 gm/cm^3 or
1 kg/liter); [Density is
mass/volume = gm/cc or kg/liter, Specific Gravity = mass relative
to the mass of the same
volume of water]
Ocean volume = 4piR2 x depth x %ocean surface
= 4 x 3.1416 x (6378)^2 x 3.73 x
0.71 = 1.354x10^9 km^3. (S x 4 x 0.7 km^3 = S x 2.8 km^3)
Density of seawater is 1.0265 kg / liter at 5
degC and salinity of 35, and therefore
mass of sea water = 1.354x 10^9 x 1.0265 x 10^12 kg = 1.39 x 10^9
x 10^12 kg = 1.39 x 10^21
(S x 4 x 0.7 x 1 x 10^12 kg = S x 2.8 x 10^12)
Salt (all principal ions) content of sea water
is .035 kg/kg of seawater,
therefore total mass of salt is:
.035 x 1.39 x 10^21 kg = 4.86 x 10^19 kg
(S x 4 x 0.7 x 1 x 10^12 x .03)kg = (S x 0.1 x 10^12)kg
and since Density of salt is c. 2.17 kg/liter (SG = 2.17)
the total volume of salt = 4.86 x 10^19 x 1 x 10^-12 / 2.17 = 2.224
x 10^7 km3
(S x 4 x 0.7 x 1 x .03 / 2 km^3 = S x 0.05 x 10^12)
If all the salt in the oceans were spread over only the continents, its thickness would be:
(2.224 x 10^7) km^3 / (1.4824 x 10^8) km^2 = 151 m.
Calculated another way using approximations, where S is the surface area of the Earth, the salt thickness would be:
Total mass of salt / Surface
area of the continents =
(Surface area of the Earth x oceanic prop x depth of oceans
x salt conc. x SG of sea water / SG salt) / (Surface area of the Earth
x continent proportion)
= (S x 0.7 x 4 x 1 x 0.033 / 2)/(S x 0.3)km = 0.154km = 154metres.
The calculated thickness is similar to the estimate
of 500 feet (170 m) of Swensen
(http://www.ci.pacifica.ca.us/NATURAL/SALTY/salty.html - if the
salt in the sea could be
removed and spread evenly over the Earth's land surface it
would form a layer more
than 500 feet thick.)
If the salt were spread over the whole of the
Earth, its thickness would be :
2.224 x 10^7 / 5.10101 x 10^8 x 1000 = 44 metres
The thickness in the oceans only would be 2.224
x 10^7 km^3 / 3.63 x 10^8 x 1000 = 61
metres.
How long would it take (residence time) to accumulate all the salt in the sea if all the sodium is derived from the weathering of continental crust?Annual volume of river water entering the oceans = 37400 (c. 4 x 10^4) km^3/yr
Total length of active ocean ridges =
56,000 km
If ocean crust is 10 km thick and spreading rate
is 3 cm/yr,
then rate of production of oceanic crust is 56000
x 10 x 3x10^-5 = 16.8 km^3/yr
(Estimated rate of production of oceanic crust
in the Mes. - Cen. = 25 km^3 /yr)
Assuming average rate of production of oceanic
crust is 20 km^3/yr,
then volume of ocean crust produced over 1.25
x 10^9 yrs = 1.25 x 10^9 x 20 km^3
If the density of oceanic crust is 2.8, then
mass of oceanic crust
= 2.5 x 10^10 x 2.8 x 10^12 kg = 7 x 10^22 kg.
Since the Earth is 4.5 billion years old, its oceans must have started to accumulate salt well before 1.25 billion years ago, and if between 1.25 and 2.5 billion years ago the oceans had accumulated the same amount of salt as at the present-day, then the same amount of salt must have been removed from the oceans as has been added since that time. Since we know the total mass of oceanic crust formed since that time (7 x 10^22 kg, assuming ocean crust production rate of 20 km^3/yr), then the sodium removed by reaction of seawater with oceanic crust must be:
(8 x 10^19 / 7 x 10^22) kg
= 1.143 x 10^-3 kg Na per kg of oceanic crust
= 0.0011 kg Na per kg of oceanic crust
= 0.11% Na
which is about 1/10th the present-day concentration of sodium in
sea water (1%).
At equilibrium the annual mass of sodium removed
from sea water would be equal to the annual mass added by the rivers
= 6.4 x 10^10 kg
Are calcium and sodium added or subtracted from seawater during hydrothermal convection in oceanic crust?Hydrothermal vent solution at c. 350C at 21N on the East Pacific Rise, ppm by weight;
If 1.7 x 10^14 kg of seawater
circulates through the oceanic crust each year (in comparison 4 x 10^16
kg of river water enters the oceans), and if it picks up 460 ppm Ca
(460 ppm = 4.60 x 10^-4 kg/kg) then:
(1.7 x 10^14 x 4.6 x 10^-4)kg
= 7.8 x 10^10 kg
of Ca is added to the seas, compared with 5 x 10^11 kg introduced
from rivers.
If the annual production of oceanic crust is 20 km^3 = 6 x 10^13 kg, then loss of Ca from the crust is 7.8 x 10^10 / 6 x 10^13 = 1.3 x 10^-3 = 0.13% Ca (Ca in oceanic basalt = 10%).
About the same amount
of Na (569 ppm; 1.7 x 10^14 x 5.69 x 10^-4 = 9.7 x
10^10) is lost anually by the oceans to oceanic basalt as a result
of the activity of circulating hydrothermal systems, an amount more than
sufficient to counterbalance the annual sodium gained from weathering (6.4
x 10^10 kg).
POTASSIUM
There is a clear removal
of K from ocean crust as a result of hydrothermal activity as indicated
by the following analyses:
vent,
seawater
ppm
ppm
K 975
380 difference = 595 ppm in favour of the
oceans
On the other hand K is clearly added to oceanic crust during the low-temperature breakdown of volcanic glass to form smectite clay minerals, in the conversion of plagioclase to sericite, and in the conversion of kaolinite to the oceanic clay mineral illite. Furthemore, potassic feldspar in continental rocks is the least amenable of the feldspar minerals to break-down by weathering. Consequently, as is indicated in the following tables, the potassium content of ocean water is relatively low compared to sodium.
Concentration
Total Amount
in the oceans
in the oceans,
kg/kg kg
Na 1.08 x 10^-2
1.42 x 10^19
K 3.8 x 10^-4
0.0502 x 10^19
Concentration of
Na and K in River Water
kg/kg
Na+ 6.30 x 10^-6
K+ 2.30 x 10^-6
CALCIUM (and CARBON)
Calcium enters the oceans as soluble calcium bicarbonate through the reaction:
anorthite + water + carbon dioxide = kaolinite + calcium bicarbonate
CaAl2Si2O8 + 3H2O + 2 CO2 (H2O+2H2CO3) = Al2Si2O5(OH)4 + Ca++ + 2HCO3-- CaAl2Si2O8 + 2H2O + CO2 (H2O+H2CO3) = Al2Si2O5(OH)4 + Ca++ + CO3--and is removed as calcium carbonate through the reaction:
Ca(HCO3)2 = CaCO3 + H2O + CO2 (H2CO3)Carbon reservoirs
Recommended
reading: Berner, R., 1999, A new look at the long-term
carbon cycle, GSA-Today, v. 9, no. 11, p. 1-6.
The conversion of anorthite
to calcite therefore involves a net reduction of CO2 from the atmo-hydrosphere
reservoir, and provides an effective negative feedback against the build
up of CO2 in the atmosphere as a result of mantle degassing. The oceans
are close to saturated in calcite and precipitation or solution of calcite
is sensitive to variations in oceanic temperature and CO2 content.
Below a depth controlled boundary named the lysocline, temperatures are
low enough and CO2 activity (decay of organic material) high enough that
biogenically produced carbonate will enter solution (see Faure, 1998, p.
142-149). In this respect it should be noted that near the partial pressure
of CO2 in near surfaces ocean water is lower than at deeper levels.
TRACE ELEMENTS IN SEA WATER
Many elements are present
in extremely small quantities in seawater. Mn is present at levels of 1-2
parts per 10 billion, whereas the Rare Earth elements Ce and Eu are present
in quantities of 1 per 10^12 and 1 per 10^14 parts, respectively. Nevertheless,
these quantities are measurable with modern analytical equipment, and they
provide a
very interesting window on the geochemical behaviour of seawater.
REE in Seawater
REE possess a number of
physical and chemical properties that make them especially useful in geochemical
studies of igneous and sedimentary rocks, ocean water/rock, and continental
rock weathering systems. All the REE are refractory with oxide condensation
temperatures similar to Sr, U and Th. They were not fractionated
during the formation of the Earth, and REE patterns are similar
to those of chondrites and the solar photosphere. The abundances quoted
above, including that for chondrite, are about 1.5 times those of chondrite
"C1" to take into account volatile loss in the formation of the Earth.
The REE content of primitive mantle (present mantle plus
crust) is further enriched 1.5 times that of the chondritic bulk
earth.
The REE exist in the
trivalent state except for Eu which may be trivalent or divalent, and Ce
which may be trivalent or quadrivalent. In the divalent state Eu is similar
to Sr and under reducing conditions enters into the same substitutions
as this element in plagioclase. Plagioclase fractionation tends therefore
to produce a negative
Eu anomaly in the residual material.
Ce may be oxidized to Ce4+
under oxidizing conditions in oceans and will precipitate from Sea water
with the formation of manganese nodules on the sea floor. Sea water therefore
exhibits a distinctive negative Ce anomaly.
In common rock forming
minerals the REE tend to substitute for Ca.
The REE are relatively
insoluble under neutral conditions of pH, and occur at high concentrations
in minor mineral phases such as monazite, sphene, allanite, apatite, and
zircon, all minerals that tend to form late in the crystallization sequence
of igneous rocks, and which usually end up in the sand-silt component of
sedimentary rocks
The Sea Water Cerium Anomaly
Although sea water contains
very low concentrations of REE, one of its most characteristic features
is the presence of a pronounced negative Ce anomaly (Ce/(2/3La+1/3Nd))
in the REE pattern.
When sea water
enters oceanic crust close to a spreading ridge it forms a hydrothermal
convection cell, wherein the descending sulphate-bearing sea water is heated
by the heat released from the crystallising basalt beneath the ridge. Iron
in the oceanic crust is converted to the ferric state by reation with the
heated sea water, which in turn progressively becomes reducing (sulphidic)
in nature. The oxidized state of the oceanic crust causes REE, and
especially Ce, to be extracted from the sea water, and the oxidized basalts
are consequently enriched in LREE and take on the -ve Ce anomaly (-1.2)
of sea water. The hydrothermal fluids emerging from the oceanic crust are
even more depleted in Ce and enriched in Eu relative to the other REE and
to the surrounding sea water reservoir. The REE's in the hydrothermal fluid
are then scavenged from the fluid by Fe oxyhydroxides which are incorporated
into sediments. These sediments will therefore tend to have accentuated
negative Ce anomalies (-1.5). As the ocean crust migrates from the ridge
however, the Fe oxyhydroxides will exchange REE with normal sea water and
will progressively develop REE patterns with Ce anomalies (-.6) less than
sea water. The overall effect therefore will be to continually increase
the negative Ce anomaly of average sea water. The fact that the -ve Ce
anomaly is almost twice as large in Pacific seawater as in Atlantic seawater
implies that hydrothermal systems are more active in the Pacific than the
Atlantic. The negative Ce anomaly of sea water also reflects the tendency
of Ce to be preferentially incorporated into manganese nodules forming
on the sea floor. (Sea water is saturated in Mn.) The initial Ce anomaly
may be related to the widespread deposition of manganese deposits during
the early Proterozoic (gondites of Africa and South America).
REE and the Cerium anomaly in deep ocean waters marginal to the East Pacific Rise.
The change in chemistry of chemical sediments has also been used to estimate the location of accumulation of Jurassic - Cretaceous Franciscan sedimentary sequences relative to the coeval oceanic ridge and continental margin.
The variation of REEs and cerium in deep ocean sediments of the paleo-Pacific.
CHEMISTRY AND THE ORIGIN OF LIFE AT BLACK SMOKERS
Russell and Hall have argued that life probably
originated about 4,200 million years ago on the bed of the primitive ocean
where extremely hot water seeping up at 150 C met cooler (c. 90 C), chemically
contrasting seawater. The sea was relatively oxidising, acidic, and rich
in iron. The acidity of the sea was caused by the high concentration of
carbon dioxide in the ancient atmosphere. The seawater provided carbonate
and phosphate ions, carbonic acid (from the dissolved atmospheric CO2),
iron, nickel and protons (H+) to the reaction. In contrast, the hydrothermal
solution
contained chemicals rich in electrons. The alkaline hot-spring fluids
carried ammonia, acetate, hydrogen sulphide ions (HS-) and molecular hydrogen
(H2) together with tungsten, organic sulphides, cyanide and acetaldehyde.
Russell and Hall believe that where the two waters
met, colloidal iron sulphide (FeS) membranes formed spontaneously and maintained
this chemical imbalance between the two waters. This membrane was semi-permeable,
and acted as a catalytic boundary where organic molecules could be synthesized.
The authors say the organic molecules that formed
along the membrane condensed to form polymers, under the influence of a
process called pyrophosphate hydrolysis. The original iron sulphide membrane
was then gradually taken over by organic material, and the first precursors
to organic cells were born.
FIGURES
Elemental Cosmic Abundances; atoms per 10^6 Si.
The Mantle - Crust - Ocean - Mantle cycle - the Earth’s geochemical and biological reservoirs.
The CO2 - weathering feedback mechanism.
The Cerium anomaly in deep ocean waters.
The variation of REEs and cerium in deep ocean sediments.
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