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W M White Geochemistry CHAPTER 15 OCEANS Cl IApTER 15 THE OCEANS AS A Cl IEMiCAl SySTEM INTRodUCTiON ntoine Lavoisier called seawater the rinsings of the Earth Given the tenuous understand ing of geological processes existing at the time the late 18th century this is a remarkably r insightful observation Most of the salts in the oceans are derived from weathering of the con tinents and delivered to the oceans by rivers But the story of seawater is more complex than this Some components of seawater are derived from hydrothermal metamorphism of the oceanic crust Other components in seawater most notably the principal anions as well as water itself are derived from neither weathering nor hydrothermal reactions These so called excess volatiles are derived from volcanic degassing Furthermore salts do not simply accumulate in seawater This point was overlooked by John lolly in his attempt to estimate the age of the Earth described in Chapter 8 from the mass of salts in the sea and the amount added annually by rivers His result 90 million years is a factor of 50 less than the actual age of the Earth The ocean is a dynamic open system and it is ul timately the balance between addition and removal of an element that dictates it concentration in the ocean This was recognized by Georg Forschhammer in 1865 when he wrote The quantity of dif ferent elements in seawater is not proportional to the quantity of elements which river water poars into the sea bat is inversely proportional to the facility with which the elements are made insoluble by general chemical or organo chemical actions in the sea One of our objectives in this chapter will be to examine the budget of dissolved matter in the oceans that is to determine the sources and sinks and the rates at which salts are added and removed from the oceans Elements also cycle between different forms within the oceans these include both organic and inor ganic solids as well as various dissolved species This internal cycling is intimately tied to the vari ous physical geological and biological processes occurring within the ocean The biota plays a par ticularly crucial role both in internal chemical cycling and in controlling the overall composition of seawater A second objective of this chapter will be to examine how elements and compounds are dis tributed within the ocean and how they cycle between various forms Lavoisier s statement also reminds usthat the oceans are part of a grander geochemical system Sediments deposited in the ocean provide a record of that system On human time scales at least the ocean appears to be very nearly in steady state It is tempting to apply Lyell s principal of uniformi tarianism and assume that the composition of the seawater has also been constant on geologic time scales There is however strong evidence that some aspects of seawater composition do change over time as we found in Chapters 8 and 9 Precisely because these variations are related to changes in other geological processes such as plate tectonics climate life and atmospheric chemistry they can tell us much about the Earth s history and the workings of the planet Interpreting these past changes begins with an understanding of how the modern ocean works and the controls on its composi tion This understanding is our main goal for this chapter SOME BACkQROUNd OCEANoanphic CONCEPTS SAliNiTy CHIORiNiTy DENSiTy ANd TEMPERATURE A useful concept in oceanography is salinity Salinity can be thought of as the total dissolved sol ids in seawater More precisely salinity is defined as the weight in grams of the dissolved inorganic matter in one kilogram of water after all the bromide and iodide have been replaced by the equiva lent amoant of chloride and all carbonate converted to oxide COz driven off This unfortunate defi Antoine Lavoisier born in France in 1843 is often called the father of modern chemistry He died at the guillotine in 1794 645 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS nition has an interesting historical basis Robert BoyleJr found that he could not reproduce total dis solved solid measurements by drying and weighing The salinity definition is due to Thorensen who would bubble Clz gas which substitutes for Br and 1 through seawater The salt could then be heated converting carbonate to oxide and a constant weight achieved Salinity is now determined by measuring electrical conductivity which increases in direct relation to the concentrations of ions in water and hence with salinity Another useful definition is chlorinity which is the hdlide concen tration in grams per kilogram measured by titration with silver and calculated as all the hdlide were chloride total halides are actually 0043 greater than chlorinity Chlorinity can also be measured by conductivity As we shall see Cl is always present in seawater as a constant proportion of total salt and therefore there is a direct relationship between chlorinity and salinity By defini tion 30 180655 Clo 151 Standard seawater which is close to average seawater has a salinity of 35000 parts per thousand ppt or o and a chlorinity of 19374 0 Open ocean water rarely has a salinity greater than 380 or less than 330 Temperature along with salinity determines the density of seawater Since density differences drive much of the flow of ocean water these are ke U 1 39 r t T r t in the oceans can be reported as potential or in situ temperature but the former is the most commonly used In situ temperature is the actual temperature of a parcel of water at depth Potential temperature denoted 9 is the temperature the water would have if brought to the surface The difference between the two is thus the temperature difference due to adiabatic expansion Since water cools when it ex pands potential temperature is always less than in situ temperature except for surface water where there is not difference The difference is small on the order of 01 C While this difference is im portant to oceanographers it is generally negligible for our purposes Temperature and salinity and therefore also density are conservative properties of seawater which is to say that they can be changed only at the surface The density of seawater is 2 to 3 percent greater than that of pure water Average seawater with a salinity of 350 and a temperature of 20 C has a density of 10247 gcc Density is usually reported as the parameter 0 which is the per mil deviation from the density of pure water 1 gcc Thus if density is 10247 gcc o is 247 Again one can distinguish between in situ and potential density po tential density being the density water would have if brought to the surface and is always lower than in situ density The difference is small a few percent and generally negligible for our purposes CiRCUlATiON of THE OCEAN ANd THE STRUCTURE of OCEAN WATER The concentrations of dissolved elements vary both vertically and horizontally in the ocean To fully understand these variations we need to know something about the circulation of the ocean This circulation like that of the atmosphere is ultimately driven by differential heating of the Earth solar energy is gained principally at low latitudes and lost at high latitudes Because the mecha nisms of surface and vertical circulation in the oceans are somewhat different it is convenient to treat them separately SURI ACE CiRCUlA TiON Surface circulation of the ocean is driven primarily by winds hence the surface circulation is some times also called the wirLd driveri circulation Figure 151 is a simplified map of the wind driven cir culation The important features are as follows I Both north and south of the climatic equator known as the Inter Tropical Convergence or lTC water moves from east to west driven by the Trade Winds These currents are known as the North Jr Robert Boyle 1627 91 was another of the founders of modern chemistry He defined the chemical element as the practical limit of chemical analysis and deduced the inverse relationship between the pressure and volume of gas a version of the ideal gas law 646 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS Figure 151 Surface and Deep Circulation of the oceans Arrows show the direction of wind driven cur rents Gray areas are regions of upwelling Red stippled areas are regions of deep water production In the Indian Ocean black arrows show current directions in northern hemisphere winter red arrows show current direction in northern hemisphere summer and South Equatorial Currents Between these two currents the Equatorial Counter Current runs from west to east I Two large gyres operate in both the Atlantic and Pacific Oceans one in the northern and one in the southern hemisphere Rotation is clockwise in the northern hemisphere and counter clockwise in the southern hemisphere The Coriolis Force an apparent force that results from the Earth s rota tion is largely responsible for this circular current pattern These currents are most intense in along the western boundaries of ocean basins a phenomenon also due to the Earth s rotation Examples of intense western boundary currents are the Gulf Stream and Kuroshio Current I The circulation in the Indian Ocean is similar but undergoes radical seasonal changes in response to the Monsoons In northern hemisphere summer the North Equatorial Current reverses and joins the equatorial countercurrent to become the Southwest Monsoon Current The Somali Current which flows to the southwest along the African Coast in northern hemisphere winter reverses direction to ow northeastward in northern hemisphere summer I Water moves from west to east in Southern Ocean the globe encircling belt of ocean south of Af rica and So America This is called the Antarctic Circumpolar Current or West Wind Drift Di rectly adjacent the Antarctic coast a counter current called the Polar Current runs east to west DENsiTy STRUCTURE ANd DEEp CIRCUIA TiON The deep circulation of the oceans is driven by density differences Seawater density is controlled by temperature and salinity so this circulation is also called the thermohuline circulation Most of the ocean is stably stratified that is each layer of water is denser than the layer above and more dense than the layer below Where this is not the case a water mass will move up or down until it reaches a level of equilibrium density Upwelling of deeper water typically occurs where winds or currents create a divergence of surface water Downwelling occurs where winds or currents produce a convergence of surface water Wind and current driven upwelling and downwelling link the surface and deep circulation of the ocean 647 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS instance we define it as 1000 m Water flows across the pycnocline only a few limited regions we can divide these into regions of intermediate water formation and deep water formation formation refers to a water mass ac quiring its temperature and salinity characteristics at the surface and sinking through the pycno cline Intermediate waters do not usually penetrate below depths of 1500 m39 deep water may pene EXAMplE 15 REplACEMENT TiME of DEED OCEAN WATER Use the simple twobox model inFigure 154 together with the followingto estimate the residence time of water inthe deep ocean Take the boundary betweenthe surface and deep water to be 1000m Assume the system is at steady state and that 14CC ratio in deep water is 10 lower than in of surface water Stuvier et al 1983 Answer Since we can assume the system is at steady state the upward and downward fluxes of both water and carbonmust be equal Denotingthe water flux by FW we may write the mass balance equation for carboninthe deep water reservoir as FWCD FWCs FCP 152 where CD CS and CF are the concentrations of carbon in deep waterand shallow water respectively and 13CP is the flux of carboncarried by sinkingparticl es The sinkingparticle flux inthusjust FCP FWCD CS 153 In other words the sinking particle flux mLst account for the difference in carbon concentration betweenthe surface and deep water We may now write a mass balance equationfor 14C indeep water by setting the loss of 14C equal to the gainof 14C 14C is lost through the upward flux of water and radioactive decay and gained by the downward flux of water and the sinkingparticle flux FWCD14CCD 7 XVDCDCACCD FWCS14CCS FCP CCs 154 where MCOD and MCCS are the 14CC ratios in deep and shallow water respectively VD is the volume of deep water and is the decay constant of 14C We have implicitly assumed that sinking particleshave the same 14C activity as surface water Subsituting153 into 154 we have FWCD CCD XVDCDCACCD FWCS14CCS FWCD 7 Cs MCCS 15 5 Rearrangingand eliminatingtermswe have 7 VD14CCD FW14CCS 7 FW14CCD 156 Another rearrangement and we arrive at VDFW 1 7 MCCD14CCS7t1ACCD14CCS 157 As we shall see later in this chapter we define steady state residence time as the amount in a reservoir divided by the flux into it or out of it Thus the above equation gives the residence time of waterinthe deep ocean notice it has units of time Substituting 01209 x 10 3 yr 1 for 7 Table 85 and 09 for 14CCDMCCs we calculate a residence time of 920 years This is somewhat longer thanthe residence time arrived at by Stuvier et al 1983 through a more sophisticated analysis We can also use this equation to calculate the average upward velocity of water Rearranging 157 we have FW VD7 14CCD14CCS1 714CCD14CCS 158 If we express the volume of the deep oceanasthe average depth d times area A we have FW Ad7 14CCD14CCS1 714CCD14CCS 159 Dividingboth sidesby A we have FWA dX14CCD14CCS1 714CCD14CCS 1510 FWA is the velocity Takingd as 3000 m we calculate FWA as 326 myr This calculation follows a similar one inBroeker and Peng 1952 649 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS trate to the bottom There are four principal regions of intermediate water production The sulfa 0mm 0 11 first is in the Mediterranean where evaporation increases salinity of surface water H to 37 380 Winter cooling further increases density causes this surface water to sink It Atlantic 501mm flows out of the Strait of Gibraltar and sinks in 013 05311 P the Atlantic to a depth of about 1000 mwhere 011 351 it spreads out This water is known as Mediterranean Intermediate Water MIW Another intermediate water known as Antarc tic Intermediate Water or AAIW is produced at the convergence at the Antarctic Polar Front at about 50 S North Pacific Intermediate Water is produced at the convergence Arctic Polar Front at about 50 N in the Pacific North Atlantic Intermediate Water is also produced at the Arctic Polar Front at 50 to 60 N Of these water masses Antarctic Intermediate Water is the densest and most voluminous There are only two regions of deep water production both at high latitudes Antarctic Bottom Wa ter AABW which is the densest and most voluminous deep water in the ocean is produced primar ily in the Weddell Sea Cold winds blowing from Antarctica cool it while freezing of sea ice in creases its salinity The other deep water mass North Atlantic Bottom Water NADW is produced around Iceland in winter when winds cause upwelling and cooling of saline MIW NADW then sinks and flows southward along the western boundary of the Atlantic In the Southern Ocean it mixes with and becomes part of the AABW Mixing between deep water and water results in a slow diffuse upward advection through the deep layer and then into the thermocline Thus whereas the flux from the surface layer to the deep one is focused the upward flux is diffuse Final return from the thermocline to the surface occurs in local ized zones of upwelling The principal upwelling zones are those along the equator where the trade winds create adivergence of surface water along the west coasts of continents where winds blowing along the coast drive the water offshore this is a process known as Ekman transport and is related to the Coriolis force and at the Antarctic divergence in the Southern Ocean With our knowledge of deep water circulation we can extend our one dimensional model Figure 153 to two dimensions Figure 154 The model illustrates several important features of the deep circulation of the oceans First no deep water is produced in either the Pacific or Indian Oceans Sec ond the Atlantic exports deep water and imports surface water Both the Indian and Pacific import deep water and export surface water Third all exchanges of deep water take place via the Southern Ocean This simple picture of deep water transport will allow us to easily understand some of the chemical differences between Pacific and Atlantic Ocean water This model can also be used to gether with 14C activities to determine the replacement time or ventilation time of deep water Stuvier et al 1983 used this model and 14C activities measured at 124 stations occupied during the GEOSECS program from 1972 to 1978 to determine deep water residence times of 275 250 and 510 years for the Atlantic Indian and Pacific Oceans respectively THE ComposiTiON of SEAWATER Figure 154 Simple two dimensional box model of ocean circulation The volumes of each reservoir are not given in units of 1018 m3 after Stuvier et al 1983 In this case the boundary between surface wa ter and deep water is taken as 1500 m Table 151 lists the concentrations and chemical form of the elements in seawater Concentrations range over 12 orders of magnitude 16 ifH and O are included From Figure 155 we see that the most abundant elements in seawater are those on the wings of the Periodic Table the alkalis the alka line earths and the halogens In the terminology we introduced in Chapter 6 these elements form hard ions that have inert gas electronic structures Bonding of these elements is predominately co valent they have relatively small electrostatic energy and large radius low Zr ratio so that in solution they are present mainly as free ions rather than complexes Elements in the interior of the 650 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS periodic table are generally present at lower concentrations These elements have higher Zr ratios form bonds ofa more covalent character and are strongly hydrolyzed The latter tendency leads to their rapid removal by adsorption on particle surfaces A few elements are exceptions to this pattern These are elements such as S Mo Tl and U that form highly soluble oxyanion complexes for exam ple SOE MoO 2 UO or soluble simple ions eg Tl Although solubility provides a guide to elemental concentrations in seawater the composition of seawater is not controlled by solubility Rather the composition of seawater is controlled by a vari ety of processes from tectonism on the planetary scale to surface adsorption desorption reactions at the atomic scale Many of the same processes that remove the elements from seawater and thus play a role in con trolling its composition also impose vertical and to a lesser degree horizontal concentration gradi ents in the ocean Table 151 also assigns each element to one of three categories based on their verti cal distribution in the water column C conservative CG conservative gas N biologically con trolled nutrient type distribution S scavenged In the following sections we will examine the behavior of each of these groups and the processes responsible for these gradients SpECiA TiON iN SEA WA TER The wide variety of elements and the relatively high concentrations of ligands in seawater leads to the formation of a variety of complexes The fraction of each element present as a given species may be calculated if the stability constants are known Chapter 6 Calculation of major ion specia tion requires an iterative procedure similar to that in Example 67 Calculation of trace element spe ciation is fairly straightforward as demonstrated in Example 152 Table 151 lists the principal EXAMplE 152 INORQANiC COMDlEXATiON of Ni iN SEAWATER Usingthe stability constants B15 for Ni complexes and the free ligand concentrationsinthe adjacent table calculate the com 18X L0 0 L0 Canon fraction of total dissolved Ni in each form Assume a NiOH 63 57 temperature of 25 C Use the following free single ion NiOHz 121 57 activity coefficients for the ligands OH 065 Cl 063 C03 NiCl 28 26 02 804 017 Use the Davies equation equation 388 to Nicos 131 45 obtaintheremainingactivity coefficients NiSO4 21 20 Answer Our first task is to calculate apparent stability constants for seawater a high ionic strength solution The ionic strength of seawater is0739 using the Davies equation we calculate a logy for Niz of 05 and logy of 0125 for singly charged species and a logy of 0 for neutral species The apparent stability constants may thenbe calculated as log 5 10g 50 log yNi Vlog yL 710g yNiL 1511 where L designates the ligand and V is its stiochiometric coefficient eg 2 for NiOHz 1 PRiNCipAl Ni COMplEXES iN SEAWATER for all others The concentration of each logy LongiLv qi form complexisgivenby Niz 1 44 NiLn BNi Lquot NiOH 353 217 0 NiOHz 812 328 0 The conservatinequat1onforN11s I Nicl 002 4124 26 2N1 N1 N10H t N10H2 NiSO4 115 O85 06 N1Cl N1304 NICOS NiC03 582 132 92 We canrewrite thisas ZNi Ni2 1 BNi2OH 5quot Ni2OH2 A little rearrangingallows us to obtainthe fractionofNi present as each species listed in the table Ni is present predominately as carbonate with minor amounts of the free ionand as chloride 651 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS species present for each element The major ions in seawater Naz K Mg Ca2 Cl SO and HCO are predominately present gt95 as free ions Many of the trace metals however are present primarily as complexes TAblE 15 CONCENTRATiONS of ElEMENTS Dissolvrd iN SEAWATER ANd RiVER WATER Average Open Ocean River Water Seawater Concentration Concentration Princi al Concentration Element pgliter uM 1 Range Dissolved Species Distribution ug liter H 11 X 108 556 X 106 H20 C 11 X 108 He 72 X 10 3 00018 He NG Li 185 265 Li C 12 Be 00023 25 X 10 5 850 pM BeOH BeOH2 SN B 461 X 103 427 BOH3 13OH 4 C 18 C 264 X 104 2200 HCO 3 CO 3 MgHCOg N Corg 1 X 102 83 various N 8540 610 N2 CG N 430 31 NOV3 N O 89 X 108 556 X 106 H20 C O 2870 179 0200 uM O2 inverse N F 1333 7013 F C 53 Ne 0164 00081 Ne CG Na 1105 X 107 4806 X 105 Na C 5300 Mg 1322 X 106 5439 X 104 M 2 MgSO 3 C 3100 A1 030 0011 1150 nM A1OH3 AlOH 1 S 50 Si 2800 100 0250 uM HASiOAy HsSiO 5 N 5000 P 62 20 035uM HPOE NaHPOMgHPOA PO 3 N 115 S 9063 X 105 2826 X 104 SO E NaSO 5 M SO j C 2840 C1 1984 X 107 5596 X 105 Cl C 4700 Ar 636 159 Ar CG K 410 X 105 1046 X 104 K C 1450 Ca 422 X 105 1054 X 104 Ca2 CaSO 3f C 14500 Sc 00006 133 X 10395 ScOH3 SN 0004 Ti 0007 14 X 10quot1 4560 pM TiOOH1 SN 10 V 178 0035 3438 nM HVO E HQVO 3 C 08 Cr 02 0004 2355 nM CrO E NaCrO 3 CrOH SN 1 Mn 002 37 X 10quot1 lt0240 nM n2MnC1 s 82 Fe 003 55 X 10quot1 005 gt6nM Fe FeCl FeOH3 SN 50 Co 0002 34 X 10395 770pM Co2 CoC1 S 02 Ni 049 84 X 10 3 312 nM Ni 2 NiC1 NiCO3 N 05 Cu 015 24 X 10 3 084nM CuCOg CuCOs 2 CuOH SN 15 Zn 038 0006 059 nM Zn ZnCl ZnSO1 N 30 Ga 00012 18 X 10 5 230 pM GaOH3 SN 009 Ge 005 00007 lt5 200 pM HAGeO4H3GeO g N 009 As 123 0016 1327 nM AsOH3 AsOH 3 SN 17 Se 0159 0002 0525 nM SeO SeO E N 0003 Br 69 X 104 863 Br C 20 Kr 032 00038 Kr CG Rb 124 145 Rb C 15 Sr 7930 905 8992 M Sr 2 C 60 Y 0017 196 X 10quot1 YCO g YCOs 5 SN 0008 Zr 0012 16 X 10quot 10300 pM ZrOH l SN 009 Nb 00046 5 X 105 N39bOgNBO5 S Mo 11 0114 MOO2 C 05 Ru lt0005 lt5 X 10 5 Rh 08 8 X 10quot1 N 652 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS TAblE 151 CONTiNUEd Average Open Ocean River Water Seawater Concentration Concentration Princi al Concentration Element ug liter M 1 Range Dissolved Species Di tribution ug liter Pd 43 X 10395 4 X 10397 0207 pM PdClE N Ag 0023 22 X 10395 130 pM AgCl N 03 Cd 0067 6 X 10quot 01000 pM CdCl CdCl2 N 002 In 000001 9 X 108 002015 pM InOH3 S Sn 47 X 10394 4 X 10396 1440 pM SnOOH g S Sb 024 0002 072 nM 3 C 0008 Te 00001 8 X 10397 0117 nM TeOHE Teog HTeO g S I 595 0468 0406 MM 0 g C 005 Xe 0065 5 X 10394 Xe CG Cs 03 23 X 10393 Cst C 0035 Ba 15 011 30130 pM Ba BaCl N 60 La 57 X 10 3 413 X 10395 450 pM LaCO 3 LaCOs 5 SN 0019 Ce 72 X 10 4 512 X 10396 28 pM OeCO g CeCO3 5 SN 0024 Pr 72 X 10 4 509 X 10396 110 pM PrCO 3 PIC03 5 SN 0005 Nd 34 X 10393 235 X 10395 340 pM NdCO g NdCOs 5 SN 0018 Sm 58 X 10 4 389 X 10396 18 pM SmCO 3 SmCO3 5 SN 0004 Eu 17 X 10 4 115 X 10396 012 pM EuCO 3 EuCOs 5 SN 0001 Gd 92 X 10 4 587 X 10396 110 pM GdCO 3 GdCOs 5 SN 0006 Tb 17 X 10 4 11 X 10396 032 pM TbCO 3 TbC03 5 SN 0001 Dy 112 X 10393 694 X 10396 1513 pM DyCO 3 DyCOs 5 SN 0005 Ho 37 X 10quot 224 X 10396 0437 pM HoCO 3 HoCOs 5 SN 0001 Er 12 X 10393 735 X 10396 1512 pM ErCO g ErCOs 5 SN 0004 Tm 2 X 10quot 121 X 10397 032 pM TmCO 3 TmC03 5 SN 0001 Yb 123 X 103 711 X 10396 1513 pM YbCO g YbC03 5 SN 0005 Lu 23 X 10394 135 X 10396 0323 pM LuCO 3 LuCOs 5 SN 0001 Hf 16 X 104 9 X 10 7 04 24 pM g SN 25 X 103 Ta 25 X 10393 14 X 10395 TaOH5 S W 0 010 54gtlt 105 W03quot C 16gtlt 10394 Re 00074 396 X 10395 ReO 5 C 00004 Os 17gtlt1T6 9X109 Ir 1 X 10396 6 X 109 Ir S Pt 5 X 105 26 X 10397 PtCl Equot PtClg C Au 2 X 10395 1 X 10397 0240 fM AuOHHQO AuCl AuCl 5 S 00001 000014 7 X 10397 022 pM HgCl E HgCl g HgC2 SN 007 T1 13 X 10 2 65 X 10 5 5878 pM C 35 X 10quot1 Pb 27 X 10393 13 X 10395 3170 pM PbCl PbC12 PbCO3 S 001 Bi 3 X 10395 14 X 10397 10500 fM BiO BiOH5 S Po At Rn Rn Fr Fr Ra 13 X 10397 58 X 103910 Ra 2 N Ac Th 2 X 105 86 X 103 50650 fM ThOH4 s 01 Pa U 33 00138 UOzC03 4 C 019 Concentrations based on single analyses or only pie1980 data are shown in italics Category C Conservative N Nu trient Biologically Controlled S Scavenged C conservative gas nonconservative gas Sources Seawater Con centrations modified from Martin and Whitfield 1983 Broecker and Peng 1982 and Quinby Hunt and Turekian 1983 and the electronic supplement to Nozaki 1997 Speciation Morel and Hering 1995 Turner and Whit eld 1981 Cantrell and Byrne 1987 Bruland 1983 Erel and Morgan 1991 River Concentrations Table 122 and modified from Martin and Whitfield 1983 Broecker and Peng 1982 653 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS CONSERVATiVE ElEMENTS The conservative elements share the property of being always found in constant proportions to one an other and to salinity in the open sea even though salinity varies All the major ions in seawater ex cept for bicarbonate are included in this group Their concentrations are listed in Table 152 This con stancy of the major ion composition of seawater which is typically expressed as a ratio to Cl is some times called the Law of Constant Proportions and has been known for nearly 2 centuries For most purposes we may state that con centrations of these elements vary in the ocean only as a result of di lution or concentration of dissolved salts by addition or loss of pure water While chemical and bio logical processes occur within the ocean do change seawater chemistry they have an insignificant ef fect on the concentrations of conservative elements The major ions do vary in certain unusual situations namely 1 in estuaries 2 in anoxic basins where sulfate is reduced 3 when freezing occurs sea ice retains more sulfate than chloride 4 in isolated basins where evaporation proceeds to the point where salts begin to precipitate and 5 as a result of hydrothermal inputs to restricted basins eg red sea brines Ca and Sr are slight exceptions to the rule in that they are inhomogeneously distributed even in the open ocean though only slightly The concentrations of these elements as well as that of HCO g vary as a result of biologi cal production of organic carbon calcium carbonate and strontium sulfateJr in the surface water and sinking of the remains of organisms into deep water Most of these biologically produced particles breakdown in deep water releasing these species into solution we explore this in greater detail be low Thus there is a particulate flux of carbon calcium and strontium from surface waters to deep waters As a result deep water is about 15 enriched in bicarbonate 1 enriched in Sr and 05 en riched in Ca relative to surface water As we shall see these biological processes also create much arger vertical variations in the concen TAblE 1 52 MAJOR IONS iN SEAWATER trations of many minor constituents Some minor and trace elements are also Figure 155 The composition of seawater The most abundant elements are those on the sides of the periodic table Elements in the interior tend to be less abundant Ion gkg ppt Percent Of present in constant proportions these in at S 35 Dlssolved SOhdS clude Rb Mo Cs Re Tl and U Vana Cl 19354 553905 dium is nearly conservative with a total 80 239649 73968 range of only about 445 All these ele HCO 0140 03941 ments share the properties that they are BOH 11 0390323 03907 not extensively utilized by the biota and Bf 00673 03919 form ions or radicals that are highly F 00013 03900 soluble and not surface reactive Nat 1077 3061 M82 1290 359 DISSOlVEd GASES ca 0412 13916 The concentrations of dissolved gases in K 0399 13910 the oceans are maintained primarily by Sr 0008 003 T A class of protozoans called Acantharia build shells of SrSO4 654 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS bonate redissolves The total amount of dissolved CO2 converted to carbonate is small compared to that converted to organic carbon However a much large fraction of biogenic carbonate sinks out of the photic zone so that the downward flux of carbon in carbonate represents about 20 of the total downward flux of carbon A larger fraction of carbonate produced is also buried so that the flux of carbon out the ocean is due primarily to carbonate sedimentation rather than organic matter sedimen tation The transport of CO2 from surface to deep water as organic matter and biogenic carbonate is called the biological pump As we might expect the biological pump produces an enrichment of C02 in the deep ocean over the shallow ocean as is illustrated in Figure 158 Many vertical profiles of ECOz show a maximum at the same depth as the oxygen minimum although the example in Figure 158 does not It occurs for the same reasons as the oxygen minimum there is more organic matter at this level and hence higher respiration and deep water is often llyounger As does oxygen enrichment the extent of enrichment of CO2 in deep water depends on the age of the water mass and the down ward flux of organic matter and therefore ultimately on the intensity of photosynthesis in the over lying water It depends additionally on the rate of calcium carbonate dissolution Biological activity also produces a variation in the isotopic composition of carbon is seawater We found in Chapter 9 that photosynthetic organisms utilize 12C in preference to 13C Thus photosyn thetic activity in the upper layer depletes surface water in 12C increasing 513C When organic matter is remineralized at depth the opposite occurs deep water in enriched in 12C Biological activity therefore imposes a gradient in SEC on the water column Figure 158 Comparing the ZCOZ with the 513C profile we see that the latter shows a pronounced maximum while the former does not Why The answer to this question is left as a problem at the end of the chapter The extent of depletion of 12C in surface water will depend on biological activity 513C will be higher in productive waters than in unproductive waters The extent of enrichment of 12C in deep wa ter as does COZ depends on the age of the deep water Old deep water will have lower 513C than young deep water SEAWATER pH ANd AlkAliNiTy We found in Chapter 6 that the pH of most natural wa ters is buffered by the carbonate system and this is cer tainly true of seawater Compared to other natural wa PH ters seawater has a relatively constant pH with a mean 7 6 7 7 7 3 7 9 3 0 3 1 3 of about 8 but the variations in dissolved COz do produce pH variations of about 03 This variation is largely due to biological activity removal of dissolved COZ by photo synthesis increases pH while release of CO2 by respira 1 02 tion decreases it The reason for this is easy to under stand At the pH of seawater bicarbonate is the pre minimu d dominant carbonate species Thus we can describe the dis E 2 solution of CO2 as COZHZOltZgtHHCO 1513 g Photosynthesis extracts COz from water so reaction 1513 is driven to the left consuming H Respiration pro 4 duces COZ driving this reaction to the right producing H For this reason the pH of the ocean decreases with depth In the profile shown in Figure 159 we see a mini 5 mum in pH at the same depth as the 02 minimum reflect ing the high rate of respiration at this depth pH is also affected by precipitation and dissolution of Figure 159 PH me e in the North Paquot calcium carbonate Since bicarbonate is the most abundant Ci c Ocean POSition 0f the oxygen miniquot carbonate species the precipitation reaction is effec mum 15 Shown tively 658 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS Ca2HCO HCaC03 1514 Here is it easy to see that precipitation of calcium carbonate decreases pH while dissolution in creases it Thus production of biogenic carbonate in surface water and its dissolution in deep water acts to reduce the vertical pH variations produced by photosynthesis and respiration Another important parameter used to describe ocean chemistry and one closely related to pH is al kalinity In Chapter 6 we defined alkalinity as the sum of the concentration in equivalents of bases that are titratable with strong acid It is a measure of acid neutralizing capacity of a solution An operational definition of total alkalinity for seawater is AlkHCO 2CO 7 BOH HZPO 2HPO i NO g OH 7H 1515 Often particularly in surface water the phosphate and nitrate terms are negligible in anoxic envi ronments we would need to include the HS ion Carbonate alkalinity is CAlk HCO 2CO 7 OH 7H 1516 which is identical to 632 One of the reasons alkalinity is important is that it can be readily de termined by titration In Chapter 6 we stated that alkalinity is conservative meaning that it cannot be changed except by the addition or removal of components It is important to understand that alkalinity is not conser vative in an oceanographic sense as is for example salinity In an oceanographic sense we define a conservative property to be one that changes only at the surface by concentration or dilution While addition and removal of components may occur through precipitation and dissolution these processes have negligible effects on conservative properties Concentration and dilution affect alka linity indeed these processes are the principal cause of variation in alkalinity alkalinity is strongly correlated with salinity However precipitation and dissolution in the ocean do signifi cantly affect alkalinity whereas the affect on salinity is negligible so alkalinity is not conserva tive in an oceanographic sense lndeed alkalinity typically varies systematically with depth be ing greater in deep water than in the surface water What causes this depth variation It might be tempting to guess that photosynthesis and respira tion are responsible However these processes have no direct effect on alkalinity When CO2 dis solves in water it dissociates to produce a proton and a bicarbonate ion In the alkalinity equation these exactly balance so there is no effect on alkalinity Production and oxidation of organic matter do affect alkalinity through the uptake and release of phosphate and nitrate but the concentration of these nutrients is generally small The main cause of the systematic variation of alkalinity in the water column is carbonate precipitation and dissolution For every mole of calcium carbonate precipi tated a mole of carbonate is removed and alkalinity increases by 2 equivalents and visa versa so the effect is quite significant CARbONATE DissolUTION ANd PRECIPITATION From the preceding sections we can see that precipitation of calcium carbonate in surface waters and its dissolution at depth is an important oceanographic phenomenon Carbonate sedimentation is also an important geological process in other respects including its role in the global carbon cycle Let s examine carbonate precipitation and dissolution in a little more detail Two forms of calcium carbonate precipitate from seawater Most carbonate shell forming organisms including the plank tonic foraminifera and coccolithophorids that account for most carbonate precipitated precipitate calcite Pteropods and many corals however precipitate aragonite even though aragonite the high pressure form of calcium carbonate is not thermodynamically stable anywhere in the ocean The sur face ocean is everywhere supersaturated with respect to both calcite and aragonite usually to depths of 1000 m or morel Nevertheless except in some rather rare and unusual situations carbonate pre 1 You might ask how aragonite can be supersaturated if it is not thermodynamically stable It is supersaturated because aragonite has a lower Gibbs Freee Energy than seawater but aragonite has a higher Gibbs Free Energy than calcite so it is unstable with respect to calcite 659 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS cipitation occurs only when biologically mediated There are two interesting questions here First why does the ocean go from supersaturated at the surface to understaturated at depth and second why doesn t calcium carbonate precipitate without biological intervention There are three reasons why the oceans become undersaturated with respect to calcium carbonate at depth First increasing FcoZ of deep water drives pH to lower levels increasing solubility This might seem counter intuitive as one might think that that increasing Fcoz should produce an increase the carbonate ion concentration and therefore drive the reaction toward precipitation However in creases in Fcoz and ZCOZ with depth produce a decrease in CO concentration This is most easily understood if we express the carbonate ion concentration as a function of Fcoz using the solubility and dissociation constants for the carbonate system equations 1221 through 1223 cog KZKIKSp7COZPCOZ 1517 H l 2 This equation shows that while carbonate is proportional to Fcoz it is inversely proportional to the square of H The pH drop resulting from production of C02 by respiration is thus dominant Car bonate ion concentrations drop by over a factor of three from the surface waters to the waters with the highest dissolved C02 The second reason is that the solubility of calcium carbonate increases with increasing pressure This results from the positive AV of the precipitation reaction Calcite and aragonite are about twice as soluble at 5000 m corresponding to apressure of 500 atms than at 1 atmosphere Third the solu bility of CaCOs changes with temperature reaching a maximum around 12 C see Example 153 As we might expect the solubility of calcite is also dependent on salinity due to the effect of ionic strength on the activity coefficients but salinity variations are not systematic with de th The kinetics of carbonate precipitation are still not fully understood in spite of several decades of research Quite a bit is known however particularly about the calcite precipitation and dissolution A number of laboratory studies eg Chou et al 1989 Zuddas and Mucci 1994 have concluded that the principal reaction mechanism of calcite precipitation in seawater is Ca CO 1 021003 1518 EXAMplE 153 PRESSURE DEDENdENCE of CAlCiTE SolubiliTy The AVrfor calcite precipitationis 37 ccmol If the apparent calcite solubility product K is430 X 10 7 molzkg2 at atmospheric pressure how will K mopk 2X107 the solubility product vary between the sea 6 g 8 surface and a depth of 5000 m Assume that AVr is 0 7 independent of pressure constant salinity a constant temperature of 2 C and that pressure increases by 01MPa for every 10 m depth in the 1000 ocean Answer The pressure dependence of the equlibrium constant is 9 10 E 2000 7 Bln K AVr 3109 f 3P g4 Integrating we have Q 3000 7 AVG P KP KP 67 4000 Sea level pressure F1 is 01 MPa the pressure at 5000 m is 50 MPa Substutiting values we can construct the graph showninFigure 1510 We see 5000 that calcite is somewhat more than twice as Figure 1510 Calculated change of the calcit soluble at a depth of 5000 mthanat the surface solubility product with depth inthe ocean 660 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS EXAMDlE 154 TEMPERATURE DEDENdENCE of CAlCiTE SolubiliTy For sea water of 350 salinity the temperature dependence of the apparent calcite solubility product may be expressed as inuni ts of moles2 per kg239 logK ABTCTD10g1 1523 where A 17834874 B 0061176 C 389469267 D71595 Mucci 1983 425 Mllero 1995 For seawater of average Ca2 ion concentration Table 152 how does the carbonate ion concentration at 42 7 which saturation is acheived vary be tween0 C and 35 C Answer The equilibrium carbonate ionconcentrationis given y co hMkg 4 1quot KCaCO3 Ca2 The concentration of Ca given in Table 152 corresponds to a molal concentration of 1028 mMkg Substituting 1523 into 4015 1424 we can construct the graph shown 0 5 10 15 20 25 30 35 inFigure 1511 Maximum solubility is 1 C achieved at about 12 C and decreases at Figure 1511 Calculated temperature dependence of the both higher and lower temperatures concentrationof carbonate ionin35o salinity seawater inequilibriumwith calcite CO 1524 t 1 In other words this simple stoichiometric expression best represents what actually occurs on an mo lecular level however other mechanisms appear to predominate at lower pH In Chapter 5 we found that the net rate of reaction can be expressed as ER EKEIL 573 If reaction 1518 is the elementary reaction describing precipitation then kCa2CO T 1519 and SK k CaCOS 1520 where k and k are the forward and reverse rate constants respectively Taking the concentration of CaCO3 in the solid is 1 the net reaction rate should be Elin k Ca2CO 7ki 1521 However Zhong and Mucci 1993 and Zuddas and Mucci 1994 found that under conditions of con stant Ca2 the overall rate equation is ingt CO T 7k 1522 where kf is an apparent rate constant incorporating both the rate constant and Ca2 concentration In other words the reaction is third order with respect to the carbonate ion rather than first order as expected This indicates that other processes must be involved and strongly influence the reaction rate Exactly what these other processes are is not yet fully understood It is known that the presence of Mg and SOE ions strongly retard calcite precipitation rates Berner 1975 Busenberg and Plummer 1985 Why is not yet fully understood One possibility is that the formation of ion pairs such as Mg2 CO and Ca2 SO E reduces the availability of reac tants Another possibility is that Mg and SO f are absorbed on the surface and thus block addition of new Ca2 and CO to the surface 661 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS In contrast to the precipitation reaction disso lution of carbonate appears to begin close to the depth where undersaturation is reached How ever dissolution is not instantaneous If we could remove the water from the ocean basins we would find that mountain peaks in the ocean are covered with white carbonate sediment but that carbonate is absent from the deeper plains and valleys This picture is very much reminis cent of what we often see in winter mountains E 2 peaks covered with snow while the valleys are bare Although the snow line the elevation 4 where we first find snow depends on tempera Q ture it does not necessarily correspond to the 00 3 C isotherm Indeed it will generally be some what lower than this The snow line is that elevation where the rate at which snow falls g5gclfng 1 just matches the rate at which it melts In the 4 ocean the carbonate snow line is called the carbonate compensation depth CCD and it is depth at which the rate at which carbonate ac 5 l l l cumulates just equals the rate at which it dis 0 0 1 0 2 0 3 0 4 solves This is alwa s lower than the de th where it begins to dissolve a depth knownpas 131550111170 Ram mgCMzyr the lysocline Like the Show line the Clepth of Figure 1512 Dissolutionrate of calcite as a func the CCD does ultimately depend on thermody tion of depth in the Pacific at 19 N 156 W as de namic factors and it varies It is deepest in the tennde by the Paterson Spheres experlment Atlantic where it as deep as 5500 m and shallowest in the North Pacific where it is as shallow as 3500 m The average depth of the CCD is about 4500 m Since aragonite is more soluble than calcite it is restricted to even shallower depths Pteropod oozes sediments composed primarily of the aragonitic shells of pteropods are largely restricted to the mid ocean ridge crests and tops of seamounts The survival of carbonate in sediment is then a question of kinetics as much as of thermodynamics The kinetics carbonate dissolution in seawater have been addressed with both laboratory and field experiments Figure 1512 shows the rate of calcite dissolution determined by an ingenious experiment performed by Peterson 1966 three decades ago Peterson hung carefully weighed spheres of calcite at various depths in the ocean for 265 days He then recovered the spheres and reweighed them to determine the rate of dissolution as a function of depth The results showed a rapid increase in disso lution at a depth of about 3500 m Since then this experiment has been duplicated several times with increasing sophistication The results show that the depth at which rapid dissolution begins the lysocline corresponds reasonably closely to the depth at which undersaturation is reached Below that depth the rate of dissolution increases rapidly and correlates with the degree of undersatura tion Laboratory studies have shown that at the pH of seawater the dominant dissolution mecha nism also appears to be the reverse of reaction 1518 eg Chou et al 1989 If that is the case then the net rate of dissolution should be SK t k kf 0013 1525 The strong dependence of the dissolution rate on carbonate ion concentration is consistent with the rapid increase in dissolution rates below the lysocline Hence field and laboratory investigations appear to yield consistent results but more experimental work on dissolution mechanisms and rates under conditions relevant to seawater is still needed 662 January 25 1998 W M White CHAPTER 15 OCEANS TLpM 0 0 100 200 39 l l l l s 28 7 O o 1 o o O E f 2 c o o g 7 Q 3 7 O o Paei it 4 r 0 Atlantic l l l l 1 Figure 1520 Profiles of dissolved and Pacific Oceans From Orians et al 1990 Ti in the Atlantic Geochemistry ditions then Fe and lln can be released from particulates rather than scavenged Both these can have higher concen trations in basins where deep circulation is so limited that deep waters become anoxic eg the Cariaco trench some Norwegian ords and the Black Sea Two vertical lln pro files are shown in Figure 1521 The maximum at 1000 1500m in the left profile is associated with the 02 minimum This profile also shows surface enrichment due to riverine input River water is less dense than seawater hence it tends to mix horizontally rather than vertically The right profile is lo cated over the Galapagos Spreading Center and shows the ef fect of hydrothermal input on Mn concentrations THE ONEDIMENSIONAI AdVECTiONiDifoSiOV ModEl Let s examine these concentration depth profiles in a bit more detail Concentration profiles such as these can be read ily modeled using a one dimensional advection aliffusion model Craig 1974 The essential assumption of such a model is that the profile observed is a steady state feature that is that the variation with depth is the same today as it was say 1000 years ago Let s begin by considering the sim ple case of the vertical variation of conservative property of ocean water such as salinity between fixed values of salinity at the top and bottom of the water column Salinity will vary only because of transport of water by our definition of conservative chemical and biological processes have no ef fect Two kinds of transport are of interest turbulent transport and vertical velocity of the water Turbulent transport is also known as lleddy diffusion lts is exactly analogous to chemical diffusion Total Dissolvalile Manganese nmolkg 05 10 0 o o o 1000 A 2000 E 0 Q 3000 o 4008 o 39 a 5000 Total Dissolvalile Manganese nmolkg 10 1 l1 Figure 1521 Profiles ofllnin the North Central Pacific left and over the Galapagos vent areas right 668 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS and may be described by the equation BC BZC W W where c is a concentration such as salinity K is the eddy diffusion coefficient and has units of mzyr and z is depth Notice that equation 1528 is identical to Fick s Second Law equation 591 except that we have replaced the chemical diffusion coefficient D with K and X with z we define 2 as be ing positive upward Adding a term for vertical velocity we have 1528 BC BZC BC m 1529 Notice that the velocity term in equation 1529 is exactly analogous to the one in equation 5159 which we derived from sediment diagenesis At steady state DcDt 0 so KB C 7 BBC 3Z2 BZ 1530 This is a second order differential equation with respect to c the solution depends on the boundary conditions These are that c is fixed at c c0 at the bottom z 0 and c cZ at the surface ZZ The solution to this equation is 02 oz 7 cofz 7 c0 1531 E 71 where 1532 exp 71 Since cz is a linear function of fz equation 1531 can be used to test the appropriateness of the one dimensional model If a truly conservative parameter such temperature of salinity is plotted against fz a straight line should result Any deviation from linearity would indicate there is sig nificant horizontal advection and that the one dimensional model is not appropriate Provided hori zontal advection is not occurring we can use equation 1531 to determine whether a particular species is conservative or not any deviation from linearity on a plot of c versus fz would indicate non con servative behavior Now let s consider a non conservative species that is actively scavenged from seawater through surface adsorption on particles We assume that the adsorption rate is proportional to the concentra tion ie first order kinetics The change of concentration with time can then be described as C K BZC C 37 W mE VC 1533 where V is the scavenging rate constant which we assume is constant with depth The sign of V is such that positive V corresponds to removal from seawater ie adsorption negative to desorption whas units ofinverse time the inverse ofwis known as the scavenging residence time and is denoted TV B2 C BC At steady state KW DE WC 1534 The solution to this equation which again depends on the boundary conditions is CZ cZFVz coGVz 1535 669 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS one of these sources or sinks might dominant and some or most of the sources might be negligible In the following sections we will discuss each of these sources and sinks Before we do however we in troduce a useful concept in marine geochemistry that of residence time RESidENCE TiME An important concept in the chemistry of seawater is that of residence time Barth 1952 Goldberg and Arrhenius 1958 Residence time T is defined as the ratio of the mass of an element in the ocean divided by the flux to the ocean ie A dAdt where A is the mass of the element of interest and dAdt is the flux to seawater Implicit in the res idence time concept is the assumption that the oceans are in steady state that is the composition does not change with time so that the flux of an element into seawater must equal the ux out of sea water Thus it does not matter whether we use the flux into or the flux out of the oceans in equation 1539 If river water is the principal source of the element equation 1539 can be re eXpressed as 1539 5 Csw X mass of seawater CSW 37 X104 CRW CRW ux of river water 15 3940 where CW and CW are the concentrations in seawater and river water respectively For water these two terms are essentially both equal to one so equation 1540 that on average a water molecule goes through the hydrologic cycle once every 37000 years For example rivers are the principal source of Na in seawater The concentration of sodium is 39 mgkg in river water and 1077 kkg in seawater so we calculate a residence time of 103 Ma for Na About half the sodium in river water is derived from cyclic salts ie it has simply been cycled through the hydrologic system If we don t count this cycling in the residence time then sodium has an ocean residence time of about 200 Ma Many ele ments have several sources and the fluxes from these are poorly known thus their residence times are Winlelown Particulates Dissolved Riverine Mn lt Particulate Nearsltore Removal Dissolved Organic EOrganic Particnlate Nearshore I 1 Mn Inorganic Inorganic Mn if 145 Part3 Mettalyfero as 75W Sedimentation Sedimentation WW ilililllim Biogenic Transport f Silica Organics to Sediments Dyj nsion out of Pelagic Sediments Figure 1523 Marine geochemistry of lLn illustrating the range of possible sources and sinks as well as internal processing of dissolved material in seawater 671 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS of fresh and sea water They may be lagoons behind barrier islands eg Famlico Sound North Caro lina river deltas eg the Rhine Delta drowned river valleys eg the Gironde Estuary France tectonic depressions eg San Francisco Bay or fjords Sannich lnlet British Columbia In the typi cal estuary there is a downstream flow of fresh water at the surface and an upstream flow of seawa ter at depth the fresh water overlies the salt water because of its lower density Depending on the geometry of the estuary the strength of the tides and the strength of the river flow these two layers will mix to varying degrees Low river ow and strong tides produce a llwell mixed estuary in which there is little vertical gradient in salinity strong river ow and weak tides lead to a 11salt wedge estuary in which a strong pycnocline develops at the interface between the layers The principal chemical changes that can occur in estuaries are as follows 0 changes in ionic stren th changes in the concentrations of major cations which affects speciation of minor components changes in pH changes in the concentration and nature of suspended matter and changes in the redox state within the water and sediment In the following paragraphs we consider these changes in greater detail As seawater mixes with river water the resulting increase in ionic strength and change in pH and solute composition induces the occulation of riverine colloids which dramatically affects trace metal chemistry As a practical matter aquatic and marine chemists often consider particulates to be the material retained on 04 mm filters and anything passing through such a filter to be lldissolved Unfortunately nature is not so neat in reality there is a continuum between truly dis solved substances and readily recognizable particles Materials between these are termed colloids Colloids are often defined as having sizes in the range of 109 to 106 m and may be separated by spe cial techniques such as ultrafiltration and membrane filtration Entities at the small end of this range would consist of 103 or fewer atoms and approach the size of larger humic acids while parti cles at the large end of the range would be retained on filter paper Colloids are a surprisingly impor tant component of natural waters Most of the Fem as well as several other readily hydrolyzed met als such as the rare earths are present as colloids of hydrous ferric oxide and metal humates rather than true dissolved species eg Boyle et al 1977 A significant fraction of humic acids is probably also colloidal Sholkovitz et al 1978 The stability of colloids depends strongly on their surface charge Because they are small colloids have high surface area to volume ratios As a result individual colloids settle very slowly and may remain in suspension indefinitely behaving essentially as dissolved species When colloids coagu late however the surface area to volume ratio decreases allowing them to settle out of suspension Whether or not they coagulate depends on the balance of forces acting between individual particles As is the case for larger particles colloids typically have a surface charge which is balanced by the ions adjacent the surface the electric double layer Chapter 6 The double layer produces a repul sive force between particles The thickness of the double layer is inversely related to the square root of the ionic strength of the solution equation 6116 Countering this repulsion is an attraction due to van der Waals interactions The strength of the van der Waals interactions decrease with the in verse square of distance and are independent of ionic strength When ionic strength increases as it does when sea water mixes with river water the diffuse outer layer or Gouy layer is compressed al lowing individual particles to approach each other more closely Once they approach within a criti cal distance the van der Waals attraction binds them together The process eventually produces par ticles large enough to settle out a process called occulation Surface charge is affected by adsorption of ions to surface sites it therefore depends on solution chemistry Surface charge is also pH dependent eg Figure 1235 Thus changes in pH and solute chemistry during estuarine mixing may also promote coagulation Eckart and Sholkovitz 1976 and Boyle et al 1977 found that occulation of humic acids occurs when Ca and Mg bind to carboxylic acid groups neutralizing their negative surface charge This allows them to coagulate and precipi tate 673 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS readsorbed or reprecipitated by l1n and Fe oxides in the upper sediment but Elderfield and Sholk ovitz 1987 nevertheless concluded that benthic flux of REE to the bay was of a magnitude similar to the riverine flux In summary chemical processes in estuaries significantly modify the riverine flux to the oceans This modification occurs primarily through solution particle reactions that occur as a result of mixing of sea and river water The results of these processes differ between estuaries because of differences in physical regime biological productivity residence time of the water in the estuary and the nature and concentration of suspended matter Some generalizations may nevertheless be made A signifi cant fraction of Fe and other particle reactive elements such as Al and the REE are removed by col loidal occulation in estuaries Though biological productivity removes nutrients SiOz N03 F04 from solution these are largely recycled within the estuary so that the net modification to the ux of these elements is probably small Chester 1991 The ux of major cations is similarly largely unaf fected Estuaries may act as a source for a number of other elements such as Ba l1n Ni and Cd through desorption and remobilization in the sediment MidiOCEAN Ridqr HydROTHERMAl SySTEMS One of the most exciting developments in geochemistry in the past 25 years has been the discovery of hydrothermal vents at mid ocean ridges Simply the sight of 350 C water black with precipi tate jetting out of the ocean bottom surrounded by a vibrant if bizarre community of organisms living in total darkness at depths of 2500 m or more was exciting But these phenomena were exciting for other reasons as well Hydrothermal systems are sites of active ore deposition so scientists were able to directly analyze the kinds of fluids that produce volcanogenic massive sulfide ores Hydrother mal activity is also an important source for some elements in the oceans and an important sink for others and has a profound effect on the composition of the oceanic crust Thus the discovery of hy drothermal vents has provided geochemists with the opportunity to put into place a major piece of the great geochemical puzzle The uids emanating at hydrothermal vents are seawater that has undergone extensive reaction at a variety of temperatures with the oceanic crust They are reduced sulfide bearing acidic and rich in dissolved metals Three kinds of venting has been observed low temperature diffuse venting in which hydrothermal solutions that have mixed extensively with seawater solutions in the subsur face diffuse slowly out of the sea oor black smokers in which high temperature usually gt300 C uid jets from sulfide chimneys and precipitate sulfide and Fe Mn oxyhydroxide smoke and white smokers in which high temperature uids ZOO 300 C jet from anhydrite chimneys and pre cipitate white anhydrite smoke THE COMPOSITION of HydROTHERMAl Fluids Samples of pure vent uids have proven difficult to obtain as vent fluids quickly mix with ambient seawater The pure vent fluid end member of the sampled mixture must therefore be calculated This is straightforward provided the concentration of at least one property of the vent fluid is known Since the temperature of the vent fluid can be determined this provides the key to calculating the vent uid end member composition The first vents discovered on the Galapagos Spreading Center were diffuse low temperature vents lt13 0 C A strong inverse correlation between Mg and tempera ture was observed and Edmond et al 1979 concluded that the pure hydrothermal fluid had a tem perature of 350 C and a Mg concentration of 0 The data from the first high temperature vents dis covered at 21 ON on the East Pacific Rise extrapolated to a similar temperature and an Mg concentra tion of 0 Figure 1530a the scatter in the data in this figure is due to the temperature probe opping about temperatures were high enough to melt the adhesive holding it in place Thus Mg appears to be quantitatively extracted from seawater in hydrothermal systems This was subsequently shown to be true of all other high temperature hydrothermal systems Having determined that the Mg concen tration of the vent fluid is 0 the concentrations of all other species in the vent uid were easily ob tained from the intercept of a plot of the concentration of the species of interest For example a plot of sulfate versus Mg extrapolates to 0 sulfate at 0 Mg Figure1530b Thus these vent fluids also have 677 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS 0 sulfate The same procedure shows the vent waters are also rich in silica and Li relative to seawa ter Figure 1531 The pure hydrothermal fluid in the first few vents discovered all had relatively homogeneous compositions and similar temperatures Subsequently discovered vents were more variable Tempera tures of uids from vents at 19 sites ranges from 220 to 403 C this range excludes diffuse ow The lowest temperatures occur where vent fluids egtltist through sediment overlying the basalt eg the Guaymas Basin in the Gulf of California Escanaba Trough on the Gorda Ridge and Middle Valley on the Juan de Fuca Ridge The highest temperatures were found in vents at the site of a recent volcanic eruption at 9 10 N on the East Pacific Rise The majority of vents have temperatures in the range of 300 C to 380 C a surprisingly narrow range This narrow range of temperatures probably reflects the large density decrease that occurs when seawater is heated beyond these temperatures Alterna tively it may reflect a sharp decrease in rock porosity beyond these temperatures L Cathles pers comm 350 22 3007 a 7 20 7 O a 7 18 7 U 250 7 16 7 7 0M o 14 E 2007 7 0 12 7 3 O E 150 7 E 10 7 A 7 E o A 8 2 1S 1007 A m o 6 7 o 507 7 4 7 0 7 0 i i i i i i f 0 10 20 30 30 M mmolk 0 10 20 40 50 W g g w Mgmmolkg 35 1000 b O 0 7 3 800 2 7 4 7 5 73 7 8 207 7 g 600 v 1 7 w 7 5 0 7 H 400 7 A o 10 l1 57 7 200 7 3900 0 l l l l l 0 0 10 20 30 40 50 U i i i i i i i i i i 439 Mgmmolkg 0 10 20 30 40 50 Figure 1530 a Temperature vs Mg concentration MgmmUlkg for several vents in the 21 N region of the EFR Figure 1531 a 5102 VS Mg for hydrothermal b MS39SUIfate PIOt for ZloN Vents From van uids from the 21 N vents b Mg Li plot for Damm at al 1985 21 N vents From Van Damm et al 1985 678 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS TAblE 1 54 COMDOSiTiON of REDRESENTATiVE HydROTHERMAl VENT Fluids 21 N Escanaba South Cleft Agtltial Volcano TAG Seawater EPR Gorda Plume Virgin Mound MAR T C 273 355 108 217 224 299 335 350 pH 33 38 54 32 44 37 39 82 Li umolkg 891 1322 1286 1718 184 845 256 6Li 66 to 10 63 to 85 323 Be nmolkg 10 13 95 383 0025 B umolkg 500 548 171 216 496 450 518 530 406 511B 300 327 101 115 342 256 268 395 C02 mmolkg 57 37 45 285 CH4 mmolkg 006 009 082 09 NH4mmolkg lt01 56 Na mmolkg 432 513 560 796 148 510 468 Al umolkg 40 52 50 53 0 015 Si mmolkg 156 195 56 69 233 135 183 0 025 HZS mmolkg 66 84 11 15 35 18 59 269 5348 14 34 78 57 73 21 Cl mmolkg 489 579 668 1087 176 559 532 K mmolkg 325 492 34 404 516 698 238 996 Ca mmolkg 117 208 334 964 102 99 105 103 llnmmollltg 67 10 001 021 359 142 659 lt004 Fe mmolkg 75 243 0 01 187 12 1 lt0006 Co nmolkg 200 lt007 Cu umolkg 15 04 164 lt0004 Zn umolkg 780 22 lt0009 Ge nmolkg 130 170 150 260 0 02 As nmolkg 30 452 13 27 Se nmolkg lt06 72 lt1 05 15 Br umolkg 802 929 1179 1832 250 847 839 Rb umolkg 27 33 80 105 37 107 145 Sr umolkg 65 97 209 312 46 51 87 87Sr 86Sr 7030 7033 07099 07028 070918 Mo nmolkg 6 115 Agnmolkg 120 lt003 Cd nmolkg 910 0 10 Sb nmolkg 18 07 13 lumolkg 99 02 05 Cs nmolkg 202 60 77 179 225 Ba umolkg 8 16 0085 Tlnmolkg 110 lt008 Pb nmolkg 1630 lt0002 From compilations of van Damm 1995 and Shanks et al 1995 EPR East Pacific Rise Mar MidAtlantic Ridge South Cleft and Axial Volcano sites are on the Juan de Fuca Ridge Concentration of sulfate in seawater Table 154 lists compositions of several representative hydrothermal vent uids all compositions are calculated on the assumption that the pure vent fluid has an Mg concentration of 0 as described above The fluids are acidic and reducing sulfide has replaced sulfate Most of the vent fluids 679 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS glass and minerals are transformed to clays such as celadonite nontronite ferric micas and smectite and oxyhydroxides In this process the rock takes up alkalis Li K Rb Cs B and U from seawater Calcite may precipitate in veins Isotopic exchange also occurs leading to higher Sr and O and lower B and Li isotope ratios in the basalt As seawater warms to temperatures around 1500 C anhydrite CaSO4 precipitates By 200 C es sentially all the Ca2 and two thirds of the SO f as well as a significant fraction of the Sr are lost in this way Addition of Ca2 to the uid by reaction with basalt will result in further removal of sulfate Thus the fluid entering the reaction zone is severely depleted in calcium and sulfate Al though anhydrite is found in altered oceanic crust it is rare Thus it is likely that much of the anhy drite precipitated in this way later dissolves when the crust cools The third major reaction in the recharge zone is loss of Mg from seawater to the oceanic crust This occurs through replacement of primary igneous minerals and glass by clay minerals such as sapolite and smectites for example by replacing plagioclase 2Mg2 4HZO ZSiO2 2CaAIZSiZO8 Z MgZCaiAlSi6020OH4 4HJr At higher temperatures chlorite forms for example by replacing pyroxene g2 6HZO 2MgFeSi206 Z MgAFeZSi4010OH8 4HJr The significance of these reactions in not only loss of Mg from the solution but also the production of H or equivalently the consumption of OH It is these reactions that account in part for the low pH of hydrothermal vent solutions This greatly increases the fluids capacity to leach and transport metals 0 Reaction Zone Seismic studies of the East Pacific Rise reveal the existence of a melt lens at shallow depth 2 km or so beneath the rise axis The depth of this magma lens is an upper limit to the depth to which hydrothermal solutions may circulate This implies pressures in the reaction zone are less than 50 MPa Geochemical observations are consistent with these geophysical constraints For instance Von Damm and Bischoff 1987 used measured SiOz concentrations in Juan de Fuca vent uids together with thermodynamic data to estimate that the fluids equilibrated with quartz at pressures of 46 48 MPa and temperatures of 390 4100 C Geothermometry performed on minerals in altered oceanic crust indicate temperatures as high as 400 5000 C By comparing a thermodynamic model of hydrothermal interactions and assuming fluids are in equilibrium with the O anh dri r o39 pi dote pyrite magnetite Seyfried and Ding 1995 estimated temperatures of 3700 to 3850 and 30 to 40 MPa for equilibration of uids from the 210 N on the EPR and the MARK area of the Mid Atlantic Ridge Under these conditions reactions would include the formation of amphiboles talc actinolite and other hydrous silicates from reactions involving ferromagnesian silicates olivines and pyroxenes the formation of epidote from plagioclase Ca2 ZHZO 3CaAIZSiZO8 Z 2Ca2AISSi3012OH ZHJr as well as the exchange of Na for Ca2 in plagioclase a process termed albitization and precipita tion of quartz 2Nafr 2CaAIZSiZO8 Z 2NaAlSi308 SiO2 2C82 The evidence for albitization comes not only from the identification of albitized plagioclase in hy drothermally altered rocks but also the inverse correlation between Na Cl and CaCl in hy drothermal fluids Figure 1533 In addition the fluid will be reduced by oxidation of ferrous iron in the rock eg 230 4H1r llFeZSiO4 2 F682 7F63304 llSiO2 ZHZO The solubility of transition metals and S increase substantially at temperatures above 350 C so sul fides in the rock are dissolved eg Clz ZHJr F633 2 HZSZiq FeClgq The metals released will be essentially completely complexed by chloride which is by far the domi nant anion in the solution as most sulfate has been removed or reduced and sulfide and CO2 will be 681 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS largely protonated at the prevailing pH Isotopic compositions of H28 in vent fluids indicates most is derived from dissolution 095 7 7 of sulfides in the rock with a smaller con tribution from reduction of seawater sul fate 0 90 o Whereas the alkalis Li K Rb Cs B 7 and U are taken up by the rock at low tem 0857 perature they are released at high tem perature Loss of K Rb and Cs begins around 150 C but loss of Li probably does 0 80 0 not begin until higher temperatures are 00 0 reached Na though is actively taken up 0 00 0 even at high temperatures by albitiza 075 O o tion Fluids make their closest approach to the magma chamber in the reaction zone 0 0 02 0 04 0 06 0 08 0 10 and magmatic volatiles may be added to fluids within this zone Hydrothermal caCl vent uids with seawater chlorinities Figure 1533 NaCl VS CaCl in hydrothermal vent have C02 concentrations as as 18 fluids Most uids define an inverse correlation that mmolkg which is substantially more results from albitization of plagioclase Only the than seawater 2 mmolkg The isotopic fluids from vents in the 9 10 N region of the EFR that composition of this carbon 513C 40 to developed after the 1991 eruption open symbols 100 is similar to that of mantle carbon deviate from this trend Composition of seawater see Chapter 9 and distinct from that of indicated by the star After Von Damm 1995 seawater bicarbonate 513C 0 Thus the excess CO2 is probably of magmatic origin Other magmatic volatiles present in the uid may include He CH4 Hz and even HZO In most cases any contribution of magmatic HZO will be insignificant compared to seawater derived HZO However uids from vents at 9 10 N on the EPR have negative 5D values Figure 1534 which suggest a small but significant contribution of magmatic water These vents developed and were sampled shortly after an eruption in 1991 Shanks et al 1995 calculated that the observed 5D values could be explained by addition of 3 magmatic water and that this wa ter could be supplied by degassing of a dike 20 km long 15 km deep and 1 m wide The magmatic wa ter would be exhausted in about 3 years The reaction zone is also the region where phase separation is most likely to occur Unlike pure wa ter which cannot boil at pressures above its critical point see Chapter 2 seawater will undergo phase separation above its critical point but the two phases produced are different from those pro duced below the critical point Below its critical point at 298 MPa and 407 C seawater boils to pro duce a low salinity vapor phase and a liquid whose salinity initially approximates that of the original liquid In the case of hydrothermal uids the vapor produced would be strongly enriched in H28 and CO2 as well as other volatiles As boiling continues the liquid becomes increasingly saline Above its critical point seawater separates into a dense brine and a fluid whose salinity initially approximates that of the original liquid As phase separation continues the fluid becomes increas ingly dilute while the brine becomes more concentrated The phase diagram F X for the system HzO NaCl shown in Figure 1535 illustrates this Seawater behaves approximately as 35 NaCl solution At a pressure of 364 MPa and 400 C a 35 NaCl solution would be above the two phase curve so only one phase exists At the same pressure and 430 C it lies just on the two phase curve and a brine containing 10 NaCl begins to separate If temperature is increased to 450 NaCl in the brine increases to 20 and decreases in the other phase to 04 Cl behaves nearly conservatively in hydrothermal solutions Phase separation and mixing be tween the uids produced by it provides the best explanation for the large variations in Cl content 100 l i i i NatCl 0 70 1 in l 682 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS i i i i down of organic matter in the sedi T ment Guaymas Basin fluids are somewhat richer in some alkalis and alkaline earths due to dissolu tion of carbonate and leaching of sediment Both the Guaymas and Escanaba fluids are transition metal poor as a result of sulfide precipitation in the sediment The Guaymas fluid is also rich in hydro carbons which are produced by thermal degradation of organic mat ter in the sediment Finally hydrothermal fluids eventually mix with seawater ei ther in the shallow subsurface or as they exit the sea oor This induces additional cooling and precipita tion Along with sulfide precipita tion mixing causes the seawater de i i i i i i i i i i rived sulfate to precipitate as an 0 0001 0x001 0 01 0 1 1 0 10 100 hydrite Precipitation of anhydrite thzrcthaCl accounts for the white smoke of Figure 1535 Pressure composition phase diagram for the Whlte smOkeFS0fWh1Ch the Vlrsm system HzO NaCl A seawater like 35 NaCl solution at M und Vent IS an example39 Preflpl39 364 MPa cross will lie above the 2 phase region at 400 C tatlon at the sea Sgrface qumkly and will thus consist of a single phase At 430 C it lies just bmlds Chlmneys Wthh can reaCh on the two phase curve and a brine of composition A begins to more than 10 meters abov the separate At 450 it lies within the 2 phase field and has Sea oor Chlmneyscon515t Prlmar39 separated into a uid of composition C and a brine of compo 11y 0f Fe and cu sul des SUCh as PY39 sition B Adapted from Bischoff and Pitzer 1989 Ute Fesz marcaSIte Fesz Pyrro39 hotite FeS chalcopyrite CuFeSz bornite Cu5FeS4 cubanite CuFezss with lesser amounts of sphalerite ZnS wurtzite ZnS galena PbS silica silicates anhydrite and barite BaSOA For the most part they are rather fragile structures subject to weathering in which anhydrite redissolves and sulfides oxidize to oxyhydroxides once venting terminates o Hydrothermal Plumes As vent fluid is diluted with seawater a hydrothermal plume is created which can rise hun dreds of meters above the vent site because of its slightly warmer temperature and therefore lower density than surrounding water Precipitation of sulfide smoke immediately above the vent re moves up to half the dissolved Fe The remained is oxidized to Fe111 and precipitated as oxyhydrox ides in the plume The half life for Fe11 oxidation in seawater is anywhere from a few minutes to a day or more depending on 02 concentrations and pH During Fe precipitation a number of elements may be coprecipitated including Mn P V Cr and As The kinetics of lln oxidation are considerably slower and Mn precipitation is generally delayed until the plume reaches neutral buoyancy and be gins to spread out horizontally Mn oxidation appears to be bacterially mediated The Fe Mn parti cles produced within the plume strongly scavenge particle reactive elements such as Th Be and the rare earths from seawater 459 400 KN c c l P75551473 MPa N O Q l 100 7 7 HydROTHERMAl HUXES The importance of mid ocean ridge hydrothermal systems in controlling the composition of seawa ter was immediately realized upon the discovery of the first hydrothermal vents Initially it ap 684 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS Pearecl that jstimagr lg the flux 0f 81839 TAblE 156 GlObAl Fluxrs TO SEAWATER ROM ments into an outo t e oceanic crust was Rid E CREST HydROTHERMAl ACTiViTy straightforward eg Edmond et al q 1989 Von Damm et al 1985 Unfortu nately the problem has proven to be not so simple A particularly important prob Hydrothermal Low T Riverine Elux Alteration Elux Elux HZ 03 15 gtlt1010 H O 13 x1010 21 gtlt1015 lem is the differences between h1 h tem 2 perature vents and diffuse vents gln dif L1 12 3399 X1010 02 1391 X1010 1394 X1010 Na 03 x1012 73 gtlt1012 fuse vents mixing between vent uids and seawater leads to extensive precipitation of the dissolved metals in the sub sea oor for such elements the global K 23 gt69 gtlt1011 1 gt 7 x1012 19 x1012 Rb 26 95 x108 19 gt 37 x108 37 x108 Cs 26 gt60 gtlt106 20 gt 38 x106 48 x106 5 6 ux depends strongly on the ratio of dif Be 30 12 X1120 12 3397 X 1012 fuse to high temperature venting Kadko Mg 1396 X 10 9 0392 X 10 12 5393 X 1013 et al 1995 Elderfield and Schultz 1996 9quot 1330 X10 quot163162 1810 Despite the uncertainties several ap Ba 24 13 X 108 1 X 1010 proaches converge on estimates of heat CH 0 67 2 4 X 1010 ux and water flux of ridge crest hy 4 39 39 11 12 drothermal activity of 2 4 X 1012 W and 2 C02 10 12 X10 11 quot2393 X 1 12 4 X 1013 kgyr respectively Elderfield 8102 46 6396 X10 8 7 X10 6394 X1 and Schultz 1996 Based on these val A1 12 6390 X110 6 X10 12 S04 84 X 10 37 X 10 ues it is possible to make rough estimates of the global ridge crest hydrothermal ux to the oceans The most recent esti mates of these uxes are given in Table HZS 085 gt96 gtlt1011 Mn 11 gt 34 x1010 28 x106 49X109 Fe 23 19 x1010 39 x109 23 x1010 Co 66 68 gtlt105 11 x108 156 Estimates of the r1ver1ne flux are 9 9 shown for comparison These uxes are all 132 13351209 X11810 lower than those estimated by Edmond et A 06 MOX 105 7392X 108 al 1979 and Von Damm et al 1985 in SS 393 9220 X104 79 X 10 many cases by an order of magnitude Even A8 7 8 11 X 106 8398 X 107 these recent estimates of uxes remain g 39 X5 39 X 6 La 41gtlt10 119gtlt10 substantially uncertain and should be 5 6 Ce 91 X 10 188 X 10 used With caution Nevertheless it ap Nd 53 105 92 106 pears reasonably well established that S 1390X105 1398X106 hydrothermal activity appears to repre E 3394105 2392105 sent a substantial flux for many elements Gd 9390 104 1397 106 including the alkalis Li K Rb Cs Be 39 X 4 39 X Dy 64 X 10 Mn Fe and Cu and an important sink for 4 others 63 Mg U Er 26 gtlt104 79 x105 Low temperature basalt seawater in Yb 1397 X 10 7396 X 105 teraction ie weatherin must also be Lu 2391X103 1399X105 39 quot 3 Pb 27 110x105 15 gtlt108 considered in assessing global uxes Es 7 6 7 timates of these uxes are included in Ta U 03918 1 6 X 10 3398 X10 3 6 X10 ble 15 6 For CO Si and Ca the low tem All numbers are in molesyear From Chen et a1 1986 Rud 39 39 z nicki and Elder eld 1993 Kadko et a1 1994 Lilley et a1 1995 Elderfield and Schultz 1996 and Staudigel et a1 1996 perature ux into the crust may exceed the high temperature ux out of it The oceanic crust is a sink for U at both high and low temperature For transition metals the low temperature ux appears to be insignificant Eor K Rb and Cs the high temperature loss from the oceanic crust appears to slightly exceed the low temperature gain for Li the high temperature loss significantly exceeds the low temperature gain Elderfield and Schultz 1996 Sr is also interesting As Table 156 shows there is no net flux of Sr to seawater from hydrothermal vents Studies of basalt similarly show the concentration of Sr in basalt 685 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS TAblE 157 Riqu HANK HydROTHERMAl does not change significantly during alteration HUXES However vent waters have 87Sr86Sr ratios between Flank Flux Rivers 0703 and 0706 generally more similar to basalt 10 10 07025 than seawater Also studies of basalts B 4119 49 X 5394 X 10 show their 87Sr86Sr ratios are increased by interac COZ 22 ng 10 11 12 tion with seawater Thus hydrothermal activity Mg 03971391 X 1 5393 X 1012 serves to buffer the Sr isotopic composition of 8102 16 1398X190 6394X1010 seawater the average 87Sr86Sr of river water is P 3392 X 119 3393 X 1012 07119 Palmer and Edmond 1989 that of seawater S 8 X10 U 3397 X1013 is 071018 This difference re ects the effect of hy Ca 20 5395 X810 1392 X1 drothermal activity We could say that the oce Ea 5 ltgt1lt106 31 X613107 anic crust is a sink for 87Sr in seawater but not for the other isotopes of Sr All numbers aIe 1 moles7981 me Kadko et 91139 Most of the heat lost by the oceanic crust occurs 1994 quoted in Elderfield and Schultz 1996 not through high temperature hydrothermal sys TAblE 158 HydROTHERMAl PlUME REMOVAl Fluxrs terns at the ridge anks Wt through lower temperature more diffuse hy Plume Hydrothermal Rivers Removal Flux Flux drothermal systems operating on 6 5 6 ridge flanks Because the flow is dif Ee X11810 30 12 X10 X11810 fuse and of low temperature these V 4393X108 5399X108 fluids have been more difficult to C 4398X107 6393X108 characterize and these systems are Mro 1399 106 xxlos not well understood Nevertheless 39 they are undoubtedly important in 8 5 8 5 09 140 X 10 7392 X 10 global uxes to and from seawater 39 Elderfield and Schultz 1996 esti 6 5 6 Ea X 186 X185 X 186 mate the water flux through ridge Nil 8398X106 5393X105 9392 X106 flank systems at 37 11 x 1015 kgyr S 2391 X106 1390 X105 1398 X106 more than 2 orders of magnitude Em 5390 X105 3394 X105 2392 X105 greater than the ridge crest flow The 611 1399 X106 9390 X104 1397 X106 most recent estimates of these fluxes D 1397 X106 6394 X104 39 X are listed in Table 157 Some of these Ey 1390106 2396104 7 9 X105 fluxes are substantial The amount of r Mg and U removed from seawater in 5 4 5 Eb X 5 X 3 X 5 this way is comparable to that re Uu 4393104 83911076 3X107 moved by ridge crest hydrothermal systems the Si flux to seawater ex ceeds that or ridge crest systems the S 1995 flux exceeds the riverine ux As we noted earlier a large fraction of the transition metals dissolved in hydrothermal fluids quickly precipitate or are scavenged by precipitation of Fe sulfides and oxyhydroxides Additional scavenging of particle reactive elements occurs through later precipitation of Mn oxides Estimates of the removal flux in hydrothermal plumes are listed in Table 158 For Be As and the rare earths the plume removal flux exceeds the primary hydrothermal flux in most cases by an order of magnitude Thus hydrothermal plumes are a net sink such elements even though they are enriched in vent uids All numbers are in moles year From compilation by Lilley et al EffECT ON THE OCEANIC CRUST Though basalt gains Mg loses silica etc such changes in major element concentrations have an in significant effect on the oceanic crust because of the high concentrations of these elements in the ba salt Minor elements may be more seriously affected Since the basalt is subducted these changes will ultimately affect the composition of the mantle Quite likely most of the water and CO2 gained by the oceanic crust during low temperature alteration are lost during subduction However there 686 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS according to Gregory and Taylor 1981 should be 04 per mil This disequilibrium reflects the effect of the Antarctic and Greenland ice sheets lce has 5180 of about 33 Storage of continental ice would tend to increase the 5180 of seawater though equilibrium crystallization of ice from seawater would result in the ice having a higher 5180 this is not the process that stores ice on continents ice is pro duced by evaporation and subsequent precipitation after IRayleigh distillation39 see Chapter 9 The half time for this process has been estimated to be about 46 Ma The half time is defined as the time required for the disequilibrium to decrease by half For example if the equilibrium value of the ocean is 0 0 and the actual value is 2 o the 5180 of the ocean should increase to 1 0 in 46 Ma It would then require another 46 Ma to bring the oceans to a 5180 of 050 etc THE ATMOSpHERiC SOURCE The atmosphere is of course the principal source and sink of dissolved gases in the ocean but it is also a surprisingly important source of other dissolved constituents as well as particulate matter in the oceans These other constituents are derived from particles in the atmosphere called aerosols Aerosols have several sources sea spray mineral dust derived from soils and desert sands volcanic eruptions condensation reactions in the atmosphere the biosphere including fires and anthro pogenic activity such as combustion of fossil fuels mining and mineral processing agriculture and the production and consumption of various chemicals Of these sources sea spray is the most important natural gas to particle conversions mineral dust and anthropogenic sources are roughly TAblE 159 ATMOS lIERIC Flux TO THE OCEANS equal in magnitude biogem39c sources are least Element Atmospheric Riverine Atmosphere important However sea spray does not rep DiSSOWBd Flux Flux RiVer resent a true flux to the oceans as it is derived N 214 X 1012 250 X 1012 086 directly from them Al 249 gtlt1011 694 gtlt1010 359 Interestingly sea spray does not have the Si 998 x 1011 667 x 1012 015 same composition as seawater Sea spray is P 959 x1009 139 x1011 007 enriched in trace metals and other sub Mn 594 X 1008 558 X 1009 011 stances This reflects the enrichment of these Fe 344 X 1010 197 X 1010 175 elements found in the surface microlayer at Ni 162 X1008 190 X1008 085 the ocean atmosphere interface Within the microlayer metals are adsorbed or com cu 13978 X 10 13960 X 10 13911 plegtlted by organic substances that form a thin zn 13955 X1007 93918 X1008 1692 lm on the sea surface As 445 X 10 133 X 10 033 The material flux from the atmosphere to Cd 489 X1006 270 X1006 181 the ocean may occur through dry deposition La 410 X1006 500 X1006 082 which includes both settling of particles C8 650 gtlt1006 630 gtlt1006 103 from the atmosphere and gas adsorption and Nd 440 X 1006 460 X 1006 096 wet deposition Wet deposition includes all Sm 910 gtlt10 15 110 gtlt1006 083 matter both particulate and gaseous first Eu 200 X1035 300 X1035 167 scavenged from the atmosphere by precipita Gd 850 X 10 140 X 1006 061 tion ie rain and snow before being deliv Dy 720 X loos 120 X1006 060 erTclli6 the 321 chemists enerall now Er 43950 X 10 800 X 10 056 agree tiat the atmosphere isgan implortant Yb 33920 X 10 13910 X1006 03929 f t f S ecies in the ocean Lu 580 gtlt1004 200 x1015 029 source or a varie y o p Pb 227 gtlt1008 100 gtlt1007 2274 quantifying the atmosphere to ocean flux is difficult There are few actual measurements A11 uxes are in moles yr Data from Duce et 611 1991 and of dry deposition rates and while there is a Greaves et al 1993 Fluxes for Cu Pb Al and Fe modi ed fair body of data on wet deposition both dry based on revised solubilities given by Chester et a1 1993 and wet deposition rates are very heteroge Mn ux estimated from the crustal Mn Fe ratio the Fe atmos neous in both space and time due to varja pheric flux of Duce and solubilities given in Chester et a1 tions in climate wind patterns and aerosol 1993 688 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS source distribution Particle concentrations can be extremely high up to 700 Mg per cubic meter of air over the Atlantic between 300 and 5 N where the Northeast trades carry mineral dust from Saharan dust storms There is also a significant ux of dust from the Asian deserts to the North Pacific though it is smaller than the Saharan dust plume In contrast particle concentrations re less than 001 ug m3 over remote areas of the South Pacific Aerosol composition also varies widely due both to difference in sources and fractionation that occurs as the aerosol is transported from the source area This variability together with the vastness of the ocean makes it difficult to derive global fluxes There is considerably more data on the composition of aerosols than on actual deposition rates Thus most estimates of the atmosphere to ocean ux are calculated by multiplying the mean atmos pheric concentration times a deposition velocity eg Chester 1990 Duce et al 1991 F CV 1541 Estimates of deposition velocity are based on models that incorporate such factors as meteorology and the nature of the sea air interface The reader is referred to Duce et al 1991 for a fuller discussion The next question we must ask is what fraction of the particulate matter deposited on the ocean sur face dissolves This depends both on the element and the nature of the particle Whereas the sea spray aerosol dissolves entirely only a fraction of the mineral aerosol dissolves Anthropogenic par ticles have intermediate solubilities For example Chester et al 1993 found that 8 and lt1 of the Al is leached by seawater from an anthropogenic aerosol collected over the UK and a mineral dust aerosol from the Arabian Sea respectively From these same two aerosols 29 and 35 respec tively of the Mn was leached Table 159 gives estimates of the atmospheric ux of dissolved matter from the oceans The esti mates were derived from the product of t J it t 39 r quot39 velocities and solu bilities These data suggest that the atmosphere is the principal source of dissolved Al Fe Cu Zn Cd Ce and Pb in the ocean However the uxes of Cu Zn Cd and Pb are primarily anthropogenic as are the uxes of Ni and As Hence these uxes do not necessarily represent the steady state for these elements The uxes of Al Fe and Ce are however primarily due to mineral dust so that even in the absence of anthropogenic activity the atmosphere may be the principal source of these ele ments SEdiMENTARy SiNkS ANd SOURCES Sediments are in one way or another the major sink for dissolved matter in the oceans There are several ways in which elements find their way into sediments 1 biologic uptake 2 scavenging by organic particles 3 scavenging or adsorption by or reaction with clay and other particles 4 pre cipitation of coprecipitation with or adsorption by hydroxides and oxides and 5 precipitation as evaporite salts In addition diffusion of dissolved species into or out of sediments may occur In the latter case sediments may serve as a source rather than a sink for a particular element BIOQENIC SEdiMENTs ANd EVAPORITES As we have seen biological activity controls the distribution of not only the major nutrients but also the micronutrients Biogenic particles incorporated in sediment are an important sink for such ele ments About 60 of the ocean floor is covered with biogenic siliceous and calcareous oozes However the cycle of biologically utilized elements is not simple For example it is estimated that the resi dence time of SiOz in seawater before biological utilization is some 200 300 years But the overall residence time is about 18000 yrs In other words a Si atom will cycle between seawater and biogenic SiOz an average of 60 to 90 times before leaving the oceans for good Recalling that Si is a refractory nutrient we can expect that labile nutrients such as P Ni and Cd are recycled many more times be fore finally leaving the system Biogenic opal and carbonate particles also scavenge other dissolved components from seawater as they fall through the water column hence biogenic sediments serve as a sink even for elements not biologically utilized Bathymetric and oceanographic factors control the distribution of biogenic sediments As we found earlier calcareous oozes are found only above the carbonate compensation depth because of the in creasing solubility of carbonate with depth The distribution of siliceous oozes on the other hand is 689 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS 80 16 0 16 80 Figure 1538 Distribution of metalliferous sediments in the oceans From Ed mond et al 1982 Thomson et al 1984 also found that when the sedimentation flux of a specific element was plotted against overall sedimentation rate a straight line resulted whose intercept was greater than zero Figure 1537 implying there was a flux of the metals even when there was no sediment flux How can this be The interpretation is as follows The red and gray clay both had the same origin on the North American continent but the red clay which accumulates slowly contained an adsorbed component of seawater Sr 87Sr86Sr0709 which could be removed by leaching Metals were also adsorbed on the gray clays but less so because they had spent a shorter had spent a shorter period exposed to seawa ter than the red clay particles The non zero intercept indicates authigenic flux of elements such as Fe and Mn the value of the authigenic flux being independent of the flux of sediment to the bottom When sedimentation rates are high the authigenic component is simply highly diluted Lack of this dilution at low sedimentation rates results in high Fe and Mn concentrations Another means of removal of elements from seawater is precipitation of oxides and hydroxides principally of lln and Fe and coprecipitation or adsorption of particle reactive elements by them This occurs in two principle ways The first is in hydrothermal plumes As we found in above hy drothermal uids are enriched in Fez and an Fe quickly oxidizes and precipitates Oxidation of llnis somewhat slower so that most precipitation is delayed until the plume becomes neutrally buoyant and begins to spread laterally As they settle out of the water column precipitated particles then scavenge other particle reactive elements from seawater When these hydrothermal particles are abundant they produce so called llmetalliferous sediment Al is not enriched in hydrothermal uids and Al in marine sediments is derived entirely from continental sources Thus sediments en riched in hydrothermally derived particles will have low ratios of AlAlFeMn Figure 1538 shows the distribution of these metal rich sediments Highest concentrations of metalliferous sedi ments are found near the mid ocean ridges particularly adjacent fast spreading regions of the EFR but the influence of hydrothermal plumes can still be seen thousands of kilometers from the ridge crest 691 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS Mn and Fe oxides and hydroxides may also precipitate directly on the seafloor with a previously existing surface acting as a nucleation site In sediment covered areas shards of volcanic glass shark s teeth and other such particles may serve as a nucleation site with the Mn Fe precipitates eventually forminga coating of up to 10 or more centimeter diameter Typically the form flattened spheres with botryoidal smooth or rough surfaces These are known as manganese nodules Solid surfaces such as the surface of a lava flow may also provide a nucleation site In this case the Mn Fe precipitates will form a coating on the surface up to several cm thick Such coatings are known as manganese crusts These nodules and crusts grow extremely slowly and occur only in areas of low sedi mentation rate They are most common in the deep basins of the Central Pacific as low sedimentation rates are most common there but they also occur in the other oceans Nodules and crusts also occur on seamounts mid ocean ridges and some areas of continental margins Mn nodules and crusts consist principally of mixtures of 5lLnOz birnessite also called 7A mangan ite toderokite also called 10A manganite and amorphous iron hydroxide FeOOH39nHZO 5MnOz birnessite and manganite all consist of primarily of sheets of MnOz but differ in their struc ture and the amount of water and other metals they contain Their average composition is given in Table 1510 As Mn and Fe particles of hydrothermal plumes nodules and crusts scavenge other parti cle reactive elements from solution so that they are usually strongly enriched in several transition metals as well as other particle reactive elements such as the rare earths as may be seen from Table 1510 This enrichment has generated interest in the possibility of mining nodules for Ni Cu and Co but while mining companies have invested in exploration and research on Mn nodules no large scale mining operations have been undertaken yet Nodules may grow by precipitation from seawater or by precipitation from sediment pore waters Nodules growing from seawater are called hydrogenous those growing from sediment pore waters TAblE 1510 AVERAQE COMDOSiTiON of MANGANESE NOdUlES ANd CRUSTS Element Average Conc Enrichment Element Average Conc Enrichment wt percent Factor ppm actor Na 194 082 B 277 277 Mg 182 078 Sr 830 22 Al 306 034 Y 310 939 Si 862 031 Zr 648 392 P 022 213 Mo 412 2747 K 064 031 Pd 00055 083 Ca 247 056 Ag 6 857 Sc 000097 044 Cd 79 395 Ti 065 114 Sn 27 135 V 0056 413 Te 50 Cr 00035 035 La 160 533 Mn 1602 1686 Yb 31 1033 Fe 1555 276 W 60 40 Co 028 1136 lr 000935 7083 Ni 048 640 Au 000248 062 Cu 026 470 Hg 050 625 Zn 0078 112 Tl 129 28666 Ga 0001 067 Pb 900 7272 B a 020 473 Bi 8 4705 From CronarL 1980 Enrichment Factor is the enrichment over average continental crust Hydrogenous nodules and crusts may grow primarily from colloidal rather than dissolved lln and Fe in seawater Strictly speaking then the process is not one of precipitation and accumulation may better term 692 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS TAblE 1511 COMDARiSON of COMposiTiON of DiffERENT MARiNE FEMN DEDOSiTS Hydrogenous Oxic Diagenetic Suboxic Diagenetic Hydrothermal Hydrothermal Crust Nodule Nodule Crust Crust Mn wt 222 3165 480 410 550 Fe wt 190 445 049 08 02 Co ppm 1300 280 35 33 39 Ni ppm 5500 10100 4400 310 180 Cu ppm 1480 4400 2000 120 50 Zn ppm 750 2500 2200 400 2020 MnFe 12 71 98 51 275 From Chester 1990 DifoSION INTO ANd OUT of SEdiMENTs Sediments can serve as a sink for dissolved matter in yet another way through diffusion of dis solved into ediment Dissolved may also diffuse out of sediment poor waters into seawater or pore water may be expelled by compaction In these latter cases sediments serve as a source of dissolved matter in seawater As we found in Chapter 5 diffusion occurs only when a com positional gradient exists Sediment pore water originates simply as seawater trapped between sedimentary particles lts composition is thus initially identical to seawater Reactions occurring within the sediment however produce changes in pore water composition establishing chemical gradients that drive diffusion into or out of the sedimentary column Furthermore as sediment is bur ied beneath subsequently accumulating material it is compacted driving pore water back into the overlying seawater producing a flux of porewater enriched components to seawater In this section we consider the examples of two elements U and Li Diffusion into sediments is an important sink for dissolved U while diffusion and porewater expulsion is an important source of dissolved Li in the oceans U is present in seawater in the VI state generally as the soluble uranyl tricarbonate species UOZC03 g4 The reduced species however U IV is relatively insoluble While seawater only rarely becomes reducing examples are the deep or bottom waters of the Black Sea some ords and the Cariaco Trench suboxic or anoxic conditions are more frequently achieved at depth in marine sediments This occurs in regions where there is a flux of high organic carbon to the seafloor as a re sult of high biological productivity in the overlying surface water Such areas occur most often on or near continental margins and cover roughly 8 of the total area of the sea floor Figure 1540a shows an example of the U profiles determined by Klinkhammer and Palmer 1991 in cores taken from the California continental shelf just south of the Monterey Fan Between 1 and 2 cm depth U pore water concentrations slightly exceed the seawater concentration 134 nmoll This re sults from release of U from labile organic phases in the upper part of the core At great depths how ever U pore water concentrations to concentrations around 5 nmoll above seawater values occur U concentrations in the coexisting solid phase increase sharply from 27 ppm to 57 ppm at depths of 6 to 9 cm Consumption of the 25 organic carbon in this core lead to suboxic conditions and reduction of U to U as well as reduction of lln and Fe Judging from the sharp increase in U concentrations in the solid phase it appears that pe values appropriate for reduction of U first occur at depths of around 6 cm Increases in pore water concentrations of lln and Fe from these cores Figure 1540b sug gests that Mn and Fe reduction begin within the top 2 or 3 cm of the core this order lLn Fe U is con sistent with thermodynamic prediction because both lln and Fe are highly insoluble in their oxi dized states Once reduced U is immobilized in the solid phase reducing it concentration in the pore waters This produces a concentration gradient that causes U to diffuse downward from seawater into the sediment 694 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS Chester R K L T Murphy E L Lin A S Berry G A Bradshaw and P A Corcoran 1993 Factors controlling the solubilities of trace metals from the non remote aerosols deposited to the sea surface by the IdryI deposition mode Mdr Chem 42 107 126 Chou L R M Garrels and R Wolast 1989 Comparative study of the kinetics and mechanisms of dissolution of carbonate minerals Chem Geol 78 269 282 Craig H 1974 A scavenging model for trace elements in the deep sea Earth Planet Sci Lett 23 149 159 Cronan D S 1980 Underwater Minerals London Academic Press Duce R A P S Liss L T Merrill E L Atlas P Buat Menard B B Hicks L M Miller et al 1991 The atmospheric input of trace species to the world ocean Global Biogeochem Cycles 5 193 259 Dymond L M Lyle B Einney D Z Piper K Murphy R Conard and N Pisias 1984 Eerromanga nese nodules from MANOP Sites H S and R Control of mineralogical and composition by multi ple accretionary processes Geochim Cosmochim Actd 48 931 949 Dzombak D A and E M M Morel 1990 Surface Complexdtion Modelling Hydrous Ferric Oxide New York Wiley lnterscience Eckert L M and E R Sholkovitz 1976 The flocculation of iron aluminum and humates from river water by electrolytes Geochim Cosmochim Actd 40 847 848 Edmond L M A Spivack B C Grant H Ming Hui C Zexiam et al Chemical dynamics of the Changjiang Yangtze Estuary Cont Shelf Res 4 17 36 1985 Edmond L M C Measures R E McDuff L H Chan R Collier et al Ridge crest hydrothermal ac tivity and the balances of the major and minor elements in the ocean the Galapagous data Earth and Planetary Science Letters 46 1979 EdmondLM Von Damm KL McDuff RE and Measures Cl 1982 Chemistry of hot springs on the East Pacific Rise and their ef uent dispersal Nature 297 187 191 Edmonds H N and L M Edmond 1995 A multicomponent mixing model for ridge crest hydrother mal fluids Earth Planet Sci Lett 134 53 67 Elderfield H N Luedtke R L McCaffrey and M Bender 1981 Benthic flux studies in Narragansett Bay Am L Sci 281 768 787 Elderfield H R Upstill Goddard and E R Sholkovitz 1990 The rare earth elements in rivers es tuaries and coastal seas Geochim Cosmochim Actd 54 971 992 Elderfield H and A Schultz 1996 Mid ocean ridge hyddrothermal uxes and the chemical compo sition of the ocean Annu Rev Earth Planet Sci 24 191 224 Elderfield H and E R Sholkovitz 1987 Rare earth elements in the pore waters of reducing near shore sediments Earth Planet Sci Lett 82 280 288 Erel Y and L L Morgan 1991 The effect of surface reactions on the relative abundances of trace met als in deep ocean water Geochim Cosmochim Actd 55 1807 1813 Eroelich P N G A Hambrick M O Andreae R A Mortlock and L M Edmond 1985 The geochem istry of inorganic germanium in natural waters L Geophys Res 90 1131 1141 Goldberg E D and G O S Arrhenius 1958 Chemistry of Pacific pelagic sediments Geochim Cos mochim Actd 13 153 212 Greaves M L P L Statham and H Elderfield 1994 Rare earth element mobilization from marine atmospheric dust into seawater Mdr Chem 46 255 260 Gregory R T and H P Taylor An oxygen isotope profile in a section of Cretaceous oceanic crust Samail Opiolite Oman Evidence for 5180 buffering of the oceans by deep gt5 km seawater hy drothermal circulation at mid ocean ridges L Geophys Res 86 2737 2755 1981 Hoyle L H Elderfield A Gledhill and M Greives The behavior of the rare earth elements during mixing of river and seawaters Geochim Cosmochim Actd 48 143 149 1984 Kadko D L Baross and L Alt 1995 The magnitude and global implications of hydrothermal ux In Sed oor Hydrothermal Systems Physical Chemical Biologich dnd Geologich Interactions Geophysich Monograph Vol 91 S E Humphris R A Zierenberg L S Mullineaux and R E Thomson ed pp 446 466 Washington AGU 697 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS Kadko D E Baker 1 Alt and I Baross 1994 Global impact of submarine hydrothermal processes RIDGE Vent Workshop 55p Keir R S 1980 The dissolution kinetics of biogenic calcium carbonates in seawater Geochim Cosmo chim Actu 44 241 252 Klinkhammer G P and M R Palmer 1991 Uranium in the oceans where it goes and why Geochim Cosmochim Actu 55 1799 1896 Kraepiel A M L I F Chiffoleau M Martin and F M M Morel 1997 Geochemistry of trace metals in the Gironde estuary Geochim Cosmochim Actu 61 1421 1436 Lilley M D F R A and I H Trefry 1995 Chemical and biochemical transformations in hy drothermal plumes In Seu oor Hydrothermul Systems Physicul Chemical Biological and Geo logicul Interactions Geophysical Monogruph Vol 91 S E Humphris R A Zierenberg L S Mul lineaux and R E Thomson ed pp 369 391 Washington AGU Mantoura R F C A Dickson and P Riely 1978 The complexation of metals with humic materi als in natural waters Estuur Coustul Shelf Sci 6 367 408 Martin I B M Kastner and H Elderfield 1991 Lithium sources in pore uids of Peru slope sedi ments and implications for oceanic fluxes Mur Geol 102 281 292 Martin I H and R M Gordon 1988 Northeast Pacific iron distributions in relation to phytoplank ton productivity Deep Seu Res 35 177 196 Martin I H K H Coale K S Johnson S E Eitzwater R M Gordon S I Tanner and e al 1994 Testing the iron phyothesis in ecosystems of the equatorial Pacific Nature 371 123 129 Measures C I B Grant M Khadem D S Lee and I M Edmond 1984 Distribution of Be Al Se and Bi in the surface waters of the western North Atlantic and Carribean Earth Planet Sci Iett 71 1 12 Millero F I 1995 Thermodynamics of the carbon dioxide system in seawater Geochim Cosmo chim Actu 59 661 678 Mucci A 1983 The solubility of calcite and aragonite in seawater of various salinities tempera tures and one atmosphere total pressure Am Sci 283 780 799 Muehlenbachs K and R Clayton Oxygen isotope geochemistry of submarine greenstones Cun Earth Sci 9 471 478 1972 Muehlenbachs K Oxygen isotope composition of the oceanic crust and its bearing on seawater Geophys Res 81 4365 4369 1976 Nozaki Y 1997 A fresh look at element distribution in the North Pacific Ocean EOS 78 221 Electronic supplement may be foun ut quothttpwwwuguorgeoselec97025ehtml Orians K I E A Boyle and K W Bruland 1990 Dissolved titanium in the open ocean Nature 348 322 325 Peterson M N A 1966 Calcite rates of dissolution in a vertical profiel in the central Pacific Sci ence 154 1542 1544 Quinby Hunt M S and K K Turekian 1983 Distribution of elements in sea water EOS 64 130 132 Redfield A C 1958 The biological control of chemical factors in the environment Am Sci 46 205 21 Rudnicki M D and H Elderfield 1993 A chemical model of the buoyant and neutrally buoyant plume above the TAG vent field Geochim Cosmochim Actu 57 2939 2957 Schaule B K and C C Patterson 1983 Perturbations of the natural lead depth profile in the Sara gasso Sea by industrial lead In Truce metuls in Seuwuter Vol C S Wong E Boyle K W Bru land I D Burton and E D Goldberg ed pp 487 503 New York Plenum Press Seyfried W E I and K Ding1995 Phase equilibria in subsea oor hydrothermal systems a review of the role of redox temperature pH and dissolved Cl on the chemistry of hot spring uids at mid ocean ridges In Seu oor Hydrothermul Systems Physicul Chemical Biological and Geological Inteructions Geophysical Monogruph Vol 91 S E Humphris R A Zierenberg L S Mullineaux and R E Thomson ed pp 248 272 Washington AGU Shanks W C I I K Bohlke and S R R II 1995 Stable isotopes in mid ocean ridge hydrothermal systems interactions between uids minerals and organisms In Seu oor Hydrothermul Systems 698 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS Physical Chemical Biological anal Geological Interactions Geophysical Monograph Vol 91 S E Humphris R A Zierenberg L S Mullineaux and R E Thomson ed pp 194 221 Washington Sholkovitz E R 1976 Flocculation of dissolved organic and inorganic matter during the mixing of river water and seawater Geochim Cosmochim Acta 40 831 845 Sholkovitz E R 1978 The occulation of dissolved Fe Mn Al Cu Ni Co and Cd during estuarine mixing Earth Planet Sci Lett 41 77 86 Sholkovitz E R 1992 Chemical evolution of rare earth elements fractionation between colloidal and solution phases of filtered river water Earth Planet Sci Iett 114 77 84 Sholkovitz E R E A Boyle and N B Price 1978 The removal of dissolved humic acids and iron during estuarine mixing Earth Planet Sci Iett 40 130 136 Sholkovitz E and D Copland 1981 The coagulation solubility and adsorption properties of Fe Mn Cu Ni Cd and Co and humic acids in river water Geochim Cosmochim Acta 45 181 189 Stanley J K and R H Byrne 1990 Inorganic speciation of zincII in seawater Geochim Cosmo chim Acta 54 753 760 Staudigel H T Plank W White and H U Schmincke 1996 Geochemical fluxes during seafloor alteration of the basaltic upper oceanic crust DSDP sites 417 and 418 In Subaluction Top to Bottom Geophysical Monograph Vol 96 G Bebout D W Scholl S H Kirby and I P Platt ed pp 19 37 Washington AGU Stumm W and I J Morgan 1996 Aquatic Chemistry New York Wiley Interscience Stuvier M P Quay and H G Ostland 1983 Abyssal water carbon 14 distribution and the age of the world oceans Science 219 849 851 Takahashi T 1989 The carbon dioxide puzzle Oceanus 32 2 22 29 Taylor S R and S M McLennan 1985 The Continental Crust Its Composition anal Evolution Oxford Blackwell Scientific Publications Thompson G A Hydrothermal fluxes in the ocean in Chemical Oceanography volS edited by I P Riley and C R p 270 337 Academic Press London 1983 Thomson L M S N Carpenter S Colley T R S Wilson H Elderfield and H Kennedy 1984 Geo chim Cosmochim Acta 481935 1948 Turner D R M Whitfield and A G Dickson 1981 The equilibrium speciation of dissolved compo nents in freshwater and seawater at 25 C and 1 atm pressure Geochim Cosmochim Acta 45 855 882 Von Damm K L 1988 Systematics of and postulated controls on submarine hydrothermal solution chemistry Geophys Res 93 4551 4562 Von Damm K L 1995 Controls on the chemistry and temporal variablility of seafloor hydrother mal systems In Sea oor Hydrothermal Systems Physical Chemical Biological anal Geological Interactions Geophysical Monograph Vol 91 S E Humphris R A Zierenberg L S Mullineaux and R E Thomson ed pp 222 247 Washington AGU Von Damm K L and I L Bischoff 1987 Chemistry of hydrothermal solutions from the southern Juan de Fuca ridge ofGeophys Res 92 11334 11346 Von Damm K L I M Edmond B Grant and C K Measures 1985 Chemistry of submarine hy drothermal solutions at 21 N East Pacific Rise Geochimica et Cosmochimica Acta 49 2197 2220 Wong C S E Boyle K W Bruland I D Burton and E D Goldberg ed 1983 Trace Metals in Seawater New York Plenum Press Zhong S and A Mucci 1993 Calcite precipitation in seawater using a constant addition technique a new overall reaction kinetic expression Geochim Cosmochim Acta 57 1409 1418 Zuddas P and A Mucci 1994 Kinetics of calcite precipitation from seawater I a classical chemical description for strong electrolyte solutions Geochim Cosmochim Acta 58 4353 4352 699 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS PROblEMS 1 In Figure 157 the 513C profile shows a pronounced maximum whereas the ECOz profile does not Why HINT Organic matter is decomposed rapidly in the ocean calcium carbonate dissolves much more slowly 2 A number of reaction mechanisms have been proposed for the precipitation of calcite from seawa ter These include Ca2 HCO g 2 H1r CaCO3 Ca CO 32 021003 Ca2 2HCO g 2 HZCO3 CaCO3 Ca2 HCO OH 2 H20 CaCO3 Suggest a series of laboratory experiments that you could performed that would enable you to dis tinguish which of these mechanisms actually occurs Assume you have a well equiped laboratory in which you can measure all macroscopic properties ie concentrations partial pressure pH carbon ate alkalinity Describe what properties you would measure and how you would use the data you obtained to descriminate between these mechanisms 3 How does calcite solubility vary with temperature and pressure in the ocean Assume that tempera ture can be represented by a simple function of pressure T Z4 e P5 l where pressure is in MPa and temperature in OC Make a plot of the solubility product as a function of depth between 0 and 5000 m using the equations in Examples 152 and 153 Hint remember to use thermodynamic temperature CHEMiCAl DATA ROM THE NORTH PAcific 4 Composition of seawater Depth Salinity Cu Ni a Calculate the molar concentrations of the major 0 054 249 90 ions in seawater listed in Table 152 b Calculate the ionic strength of this solution 75 3398 03969 2399 88 185 3392 091 379 84 c Using the equilibrium constants in Table 121 calculate the concentration of carbonate ion in equi 375 343905 13945 53926 70 librium with this solution at 25 C 595 34 19 749 60 d Calculate the total alkalinity of this solution 780 3419 215 907 52 assuming a pH of 81 985 3437 238 964 48 1505 3455 28 979 45 5 Use the one dimensional udvection di usian model 2025 3461 318 106 47 in the depth interval of 595m to 4875 m and the 2570 3455 346 103 50 chemical data from the North Pacific in the 3055 3466 39 109 54 adjacent table to answer the following questions 3533 3456 426 10 63 For this locality the ratio K was determined to be 4000 3467 457 108 66 2300 and 0 to be 4 Is salinity conservative and the 4635 3468 503 103 74 one dimensional model applicable Make a plot of S vs fz equation 1533 Are Cu Ni and Al con 4875 343968 53934 10394 79 servative Are they being produced or scavenged concentrations in anIkg salinityin PPt For each of these elements find a value of V or to fit the one dimensional advection diffusion model to the data 6 Stanley and Byrne 1990 give the following stability constants for Zn complexes in seawater Zn2 Cl 2 ZnCl 5 7040 Zn2 HCO g 2 ZnHCOgr 5 083 700 January 25 1998 W M White Geochemistry CHAPTER 15 OCEANS Zn cog 2 ZnC03 5 287 Zn 200 2 ZnCO3 5 441 an H20 2 ZnOH H 5 925 Zn so 2 cho 5 090 Using these stability constants a pH of 81 the ligand concentrations given in Table 152 and the equilibrium concentration of carbonate ion calculated with the equilibrium constant in Table 61 for a temperature of 25 C calculate the fraction of Zn present as each of these species plus free an 7 Using the flows of water through mid ocean ridge crest and flank hydrothermal systems and the mass of the oceans given in Appendix 1 how long does it take to cycle the entire ocean through these systems 8 In the San Clemente Basin off the southern California coast Barnes and Cochran 1990 found that U concentrations in sediment pore waters decreased to 335 nMl in the top 7 cm of sediment Assuming an effective diffusion coefficient corrected for porosity and tortuosity in the sediment of 681 cmzyr calculate the flux in nMcmz of U from seawater to sediment in this locality 9 Using the fluxes of U to the ocean in Table 1511 estimate the residence time of U in the ocean 701 January 25 1998 W M White Geochemistry APPENDIX llll UMMAPY OF IMPORTANT EQUATIONS EQUATiONs of STATE Ideal GasLaw PV NRT Coefficient of Thermal Expansion Compressibility Van der Waals Equation RT a P 7 V7 b 2 THE LAWS of THERMdyNAMiCS First Law AU Q W 1 written in differential form 1U dQ 1W 2 work done on the system and heat added to the system are positive The first law states the equivalence of heat and work and the conservation of energy Second Law 1 Q rev T d S 3 Two ways of stating the second law are Every system left to itself will on average change to a condition of maximum probability and Heat cannot be extracted from a body and turned entirely into work Third Law lim S 0 T gt 0 3 This follows from the facts that S R ln 9 and Q 1 at T 0 for a perfectly crystalline pure substance PRiMARy VARiAblES of THERMOdyNAMiCS The leading thermodynamic properties of a fluid are determined by the relations which exist between the volume pressure termperature energy and entropy of a given mass of uid in a state of thermodynamic equilibrium I W Gibbs The primary variables of thermodynamics are P V T U and S Other thermodynamic functions can be stated in terms of these variables For various combination of these variables there are W M White Geochemistry APPENDIX Illl EQUATION UMMARY characteristics functions The characteristic function for S and V is one of the primary variables U Thus 1U TdS PdV 5 OTHER IMPORTANT THERMOdyNAMiC FUNCTiONS What then is the use of thermodynamic equations They are useful precisely because some quantities are easier to measure than others M L McGlashan Enthalpy H E U PV 6 In differential form in terms of its characteristic variables dH TdS VdP 7 Helmholtz Free Energy A E U TS and dA E PdV SdT 9 Gibbs FreeEnergy G E H TS 10 The Gibbs Free Energy change of a reaction at constant temperature and pressure is AGr AHr TASr 10a and dG VdP SdT 11 Your choice of which of these functions to use should depend on What the independent variables in your system are In geochemistry P and T are the most common independent variables so the Gibbs Free Energy is often the function of choice EXACT DiffERENTiAlS ANd TIIE MAXWEll RElATiONS Any expression that may be written MXydxNXydy 12 is an exact differential if there exists a function z fxy such that fXy MXydXNXydy 13 The total differential of the function zxy is written B Bz dz Z dxi dyMdXNdy 14 BX y By X If dz is an exact differential then BM BN 7 7 15 By BX which is equivalent to 6 1 2 16 W M White Geochemistry APPENDIX ITIT EQUATION UMMARY All thermodynamic variables of state are exact di erentials Thus the practical application of the properties of exact differentials can be illustrated as follows Equation 11 dG VdP SdT has the form dz MXydXNXydy since V and S are functions of temperature and pressure Equation 11 may also be written as 8G dG 7 dP 8P T and comparing equations 11 and 13 we conclude that dT 17 18 19 Applying the rule embodied in Equation 15 we can conclude that al is 20 aT P aP T Playing similar games with Equations 5 through 9 we can develop a series of relationships 3T 3P from dE 7 7 7 21 av s as V f dH 3T 7 9V 22 rom 7 7 7 aP V as P P as from dA 7 23 Equations 20 23 are known as the Maxwell Relations DERiVATiVES of ENTROpy as pressure 706V 24 T as Cv as 012 temperature 7 and 7 7 25 26 T V T 3T P T 1 as 7 0c 27 V0 ume 7 7 7 3 B T DERiVATiVES of ENTHAlpy pressure Vl 7 OCT 28 T t t 3H C 29 empera ure W 7 P P DERiVATiVES of ENEqu BU BU temperature C and C 7POCV 30 31 aT V v aT P P W M White Geochemistry APPENDIX Illl EQUATION UMMARY volume T7606 7 P 32 DiffERENCE bETWEEN CP ANd Cv 33 THE Gibbs PHASE RUlE The Gibbs Phase Rule is a rule for determining the degrees offreedom of a system f c p 2 34 f is the number of degrees of freedom c is the number of components and p is the number of phases The minimum number of components needed to describe a system is c N R THE ClApEyRON EQUATiON The slope of a phase boundary in PT space is Q A V r dP ASr SolUTiONS Raoult s Law applies to ideal solutions Where N is the number of species and R is the number of reactions possible between these species 36 Henry s Law applies to very dilute solutions and state that the partial pressure of a component in solution is proportional to it mole fraction Pi hXi for Xi ltlt 1 CHEMiCAl POTENTiAl 37 Chemical potential is defined as Where nj is the number of moles of the id1 component In multicomponent systems the full expression for the Gibbs Free Energy is dG VdP SdT 2 Midni i W M White Geochemistry APPENDIX Illl EQUATION UMMARY THE GibbSrDUHEM RElATiON At equilibrium and at constant pressure and temperature 40 THERMOdyNAMiC VARiAblES iN IdEAl SolUTiONS 0 Mi ideal Hi RT 111 Xi 41 Adeealmjxing 0 and therefore Videal Z XiVi Z Xivi i i Adeealmixing 0 and therefore ideal Z Xihi Z Xi i i i AS ideal mixing 39RZ Xiln Xi i Sideal solution Xisi39RZ Xiln X 42 1 1 AGideai mixing RTE X11 X1 43 1 O Gideal solution XiHi RTE X11 X 44 1 1 THERMOdyNAMiC VARiAblES iN NOerdEAl SolUTiONS Fugacity Fugacity can be thought of as the escaping tendency of a gas in nonideal solutions Because systems tend toward ideal at low pressure it has the property iiimo Pi 1 45 7 0 fi and Hi 7 Mi RT In F 46 1 Activity Activity is defined as hence W M White Geochemistry APPENDIX Illl EQUATION UMMARY The activity in an ideal solution is aiideal Xi 49 The activity coefficient 7 is defined as 81 X1 11 50 When Henry39s Law law holds Xi hi 51 The DebyeH ckel equation is used to calculate activity coefficients in aqueous solutions It is 2 Azi 10 g Y1 1 B Where z is charge I is ionic strength a is the hydrated ionic radius significantly larger than ionic radius and A and B are solvent parameters I is calculated as 52 7 1 2 I 7 E mizi 53 1 Excess Free Energy and activity coefficients G RT 2 xm 54 EXCESS 1 1 1 Excess Free Energy and Margules Parameters of a Regular Solution Gex XIXZWG 55 Excess Free Energy and Margules Parameters of an Asymmetric Solution Gexcess WG1X2 WGZX1X1X2 56 EQUilibRiUM CONSTANT The equilibrium constant is defined as 3P RT 57 It is related to the Gibbs Free Energy change of the reaction by K AG RT 5 8 It is related to enthalphy and entropy changes of the reaction by AH Asf 1 K 7 59 n RT R Pressure and temperature dependencies of the equilibrium constant are 0 am K AV 7 60 T OXidATiON ANd REdUCTiON The redox potential is related to the Gibbs Free Energy change of reaction as AG nFE 61 W M White Geochemistry APPENDIX llll EQUATION UMMARY Where E is the redox potential n is the number of electrons exchanged and F is the Faraday constant The Nernst Equation is EEoi 1n Ha i HF 62 The p8 is defined as 63 and is related to hydrogen scale redox potential EH as 64 KiNETiCS Reaction Rates For a reaction such as aA bB ltgt cC dD A general form for the rate of a reaction is 1dA1dB 1dC 1dD nAnB nc nD 65 adtbdt cdt ddt kA B C D Where nA etc can be any number For an elementary reaction this reduces to 1dA1dB 1dC 1dD ab 66 a at b at c at quota alt kA B The temperature dependence of the rate constant is given by the Arrhenius Relation 67 68 69 70 W M White Geochemistry EQUATION UMMARY APPENDIX llll The temperature dependence of the diffusion coefficient is Diagentic Equation TRACE ElEMENTS Equilibrium or Batch Partial Melting Fractional Partial Melting i iDR D 1e Zone Refining 7 Equilibrium Crystallization Fractional Crystallization ISOTODE GEOCHEMiSTRy Binding Energy per Nucleon Eb Basic Equation of Radioactive Decay Isotope Growth or Isochron Equation R R0 RPD 6 i 1 71 72 73 74 75 76 77 78 79 80 W M White Geochemistry APPENDIX IIIIII 0ME MATHEMATICS IUSEFUL IN GEOCHEMISIRY LiNEAR REQRESSiON Fitting a line to a series of data is generally done with a statistical technique called least squares re gression Real data are not likely to fall exactly on a straight line each point will deviate from the line somewhat The idea of least squares regression is to find the best line fitting the data by minimizing the squares of the devialtliohs frolrln the regression line The quantity to be minimized is iZlez7i21y7a7bx2 1 This is know as the sum of the squares of the deviations from the line y a bx The use of the squares of the deviations means that large deviations will affect the calculated slope more than small deviations By differentiating equation 1 it can be shown that the minimum value for the left side occurs when the slope is 2 x17 m 7 a b 7 2 2 XI 7 3 02 where X and y are the means of x and y respectively and xi and yi are the ith pair of observations of x and y respectively We can see from 723 that the regression slope is the product of the deviations of x and y from the mean divided by the square of the deviations of X from the mean A more convenient computational form of 2 is 7 2 xi X 7 Vin b 7 3 2 7 izn The intercept is then given by a y bX 4 Because real data never fit a line exactly it is of interest to know the error on the estimate of slope and intercept The error on the slope is given by 7 211 2X13i7 n2 1 ob W y lIn72IIZX3722nIl 5 The error on the intercept is 2 2 2 XiYi WHY 1 5amp2 1 Ga ZX YH 2 2 3 2 2 6 2X17xn 2X17xn n Statistics books generally give an equation for linear least squares regression in terms of one depen dent and one independent variable The independent variable is assumed to be known absolutely With geochemical data both x and y are often measured parameters and have some error associated with them These must be taken into account for a proper estimate of the slope and the errors associated with it In some cases the errors in measurement of x and y can be correlated and this must also be taken into account The socalled twoerror regression algorithm does this This is however considerably less straightforward than the above The approach is to weight each observation according to the measurement error the weighting factor will be inversely proportional to the analytical error A solution written in the context of geochronology has been published by York 1969
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