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Earth System Science II - The Oceans as a Chemical Reservoir.

     In the last lecture I discussed the processes by which material is transferred from the Earth's mantle to the Crust and back again to the mantle.  This  week I want to explain how that differentiation effects specific elements.

     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, C
Elemental Cosmic Abundances; atoms per 10^6 Si.

Periodic Table.

       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.


        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
    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

       During 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.

                        What happens to the Ca, Na, and K released to the oceans as bicarbonate, 
and what are the quantities involved?

        The Ca, Na, and K released into solution during weathering is transported by rivers to the oceans representing a primary cyclic reservoir. However, the alkali elements sodium and potassium are also removed from the oceans into secondary cyclic reservoirs or permanently returned to the mantle via reaction with oceanic crust. Ca is removed to
form  limestone or buried as hydrocarbons as part of a Carbon sub-cycle or to marine biological sub-cycles, and Na and K to evaporite reservoirs and the oceanic crust, respectively.

              THE DATA:
        Concentration   Total Amount
        in the oceans   in the oceans,
          kg/kg           kg
Ca       0.00041        0.0545 x10^19   Ca2+
Na       0.01077        1.42   x10^19   Na+
K        0.00038        0.0502 x10^19   K+
Cl       0.0195         2.57   x10^19   Cl-

    % contribution of various sources to Na content of river water (Berner and Berner,
    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,
(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
( - 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

        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
                                                               = 4 x 10^16 kg
    Total sodium entering the oceans = 4 x 10^16 x 7.2 x 10^-6 = 2.88 x 10^11 kg
                                                            = 288 x 10^9 kg
                                                            = 288 x 10^12 grams
                                                            = 288 Tg
    Total Na/yr entering oceans due to weathering = 4 x 10^16 x 1.6 x 10^-6 = 6.4 x 10^10 kg
    Total Salt in Seawater =  4 x 10^19 kg
    and if salt in evaporites =  4 x 10^19 kg
    Residence time of Na = 8 x 10^19 / 6.4 x 10^10 = 1.25 x 10^9 (1.25 billion years)
    (Residence time of Na is the time is would take to add the present mass of sodium in seawater if the only source was the `weathering' sodium brought to the oceans by rivers.)

    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;
   (pH of the vent solution is 4.0 whereas normal seawater is about 8.)
                         vent                               seawater
             ppm                                   ppm
Cl       17300 =   0.0173 kg/kg    19500  = 0.0195 kg/kg
Na       9931  =   0.0099 kg/kg    10500  = 0.0105 kg/kg
Mg         -                                       1290  = 0.00129 kg/kg
SO42-    -                                         905
H2S      210                                        -
Ca        860   = 0.00086 kg/kg      400     = 0.0004 kg/kg
K          975                                    380
Sr            8                                         8
Si         600                                        3
Ba        5-13                                      2x10^-2
Zn          7                                          5x10^-3
Mn       33                                         1x10^-4
Fe        101                                        2x10^-4

        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).


        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
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.


        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.


    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.


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.

Carbon reservoirs.

The Cerium anomaly in deep ocean waters.

The variation of REEs and cerium in deep ocean sediments.

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The Cerium anomaly in deep ocean waters.

The variation of REEs and cerium in deep ocean sediments.

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