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Geochemistry of the upper part of the solid Earth: Lecture 1: The geochemical toolbox and first-order implications for major earth reservoirs Lecture 2: The geochemical toolbox (continued). Brief introduction to the continental crust Guest lecture: Bruno Dhuime (crustal evolution) Lecture 3: Mid-ocean ridge magmatism Lecture 4: Intraplate magmatism Lecture 5: Melt generation in subduction zones Lecture 6: The subcontinental mantle lithosphere The Geochemical Toolbox and first-order implications for major Earth reservoirs • Major elements Bulk Earth composition and how we know it Major element composition of Earth reservoirs • Trace elements Factors controlling trace element partitioning How these factors determine trace element distribution among Earth reservoirs • Isotopes Radiogenic isotopes Stable (conventional and non-conventional) isotope systems "Mass-independant" fractionation Goldschmidt's classification of Elemental Behavior Victor Goldschmidt (1888-1946) White (1997) Bulk Silicate Earth (BSE) Bulk Earth CaAl Ni other Fe Na, K, Al other Ca O Fe Mg Mg O Si Si Ni First order differentiation of the Earth other Fe Si? Core How do we know the bulk Earth composition? Core: From geophysics, especially seismic and mineral physics data, and from analogy with iron meteorites. Bulk silicate Earth (BSE): From intersection between partial melting trend (upper mantle peridotites) and fractional condensation trend (chondritic meteorites). Assumptions: • Upper mantle peridotites are representative of whole mantle Ma ntle p erid otite • Trends obtained from chondritic meteorites are applicable to the Earth te e M s te ori BSE s Jagoutz et al. (1979) Another way to estimate BSE composition Constant chondritic Sm/Yb ratio Peridotite melt extraction trend McDonough & Sun (1995) Estimated BSE TiO2 content Though different in detail, this method is also based on comparing peridotite melt extraction trend with chondritic meteorite trend, so same assumptions hold. Condensation temperatures in the solar nebula Albarede (2003) Abundances of refractory elements are ~ 2.75 times higher in the BSE than in C1 chondrites. Ratios of refractory elements vary little in chondrites. Traditionally Earth has been assumed to have same refractory element ratios as chondrites. But this view has changed in the past decade with very large consequences for our understanding of crust-mantle evolution. More on that later... Continental crust Continental crust extraction overall composition: andesite Na K BSE (or "primitive mantle") overall composition: lherzolite Fe Na, K, Ca Al other other Al Ca Fe O Mg Si Mg O Si Residual mantle ? Composition depends on fraction of mantle from which crust extracted m = mass C = concentration of element x = mass fraction of crust cr = crust rm = residual mantle BSE = Bulk Silicate Earth Mass Balance mBSE = mcr + mrm CBSE*mBSE = Ccr*mcr + Crm*mrm CBSE = Ccr*mcr/mBSE + Crm*(mBSE-mcr)/mBSE CBSE = Ccr*x + Crm*(1 - x) Crm = (CBSE - Ccr*x)/(1-x) mcr = 2.20 x 1022 kg mBSE = 4.002 x 1024 kg (whole mantle) x = 0.0055 mBSE = 1.96 x 1024 kg (upper mantle) x = 0.020 Accuracy of calculated concentrations in residual mantle depends on: 1) Estimated Primitive mantle concentrations (CBSE) 2) Knowledge of crustal concentrations (Ccr) 3) Fraction of mantle from which crust is extracted Concentrations normalized to BSE Effect of crustal extraction on major element composition of mantle continental crust residual whole mantle residual upper mantle Two endmember cases considered: crustal extraction from whole mantle and crustal extraction uniquely from upper mantle For most major elements, effect of crustal extraction on mantle composition is minor. For K, a highly incompatible element, crustal extraction has a large effect, especially if crust extracted only from upper mantle. White (1997) Factors controlling partitioning during melting: ionic radius and charge LIL REE HFSE White (1997) Contours: clinopyroxene/liquid partition coefficients (Kd), where Kd = Concentration in phase/concentration in liquid LIL= Large Ion Lithophile REE = Rare Earth Elements HFSE = High Field Strength Elements LIL vs. HFSE Large Ion Lithophile elements: Large radii and low charge (low z/r) highly incompatible during magmatic processes and highly mobile in hydrothermal fluids. High Field Strength Elements: Small radii and high charge (high z/r) form strong bonds that are difficult to break and are therefore immobile in hydrothermal fluids. Provides a first order explanation for why LIL and HFSE show similar behavior in mid-ocean ridge magmas (dependant essentially on magmatic processes) and divergent behavior in arc magmas (dependant on both magmatic and hydrothermal processes). Rare Earth Elements (REE) - Ionic radius 3+ charge White (1997) Explains systematic behavior of REE in most minerals and Eu anomalies found in certain rocks REE Kd values for major mantle phases In most mantle phases, light REE (LREE) are much more incompatible (low kd) than heavy REE. Positive Eu anomaly in plagioclase reflects substitution of Eu2+ for Ca2+ White (1997) Effect of partial melting on REE spectra Albarede (2003) Because LREE are more incompatible than HREE in most mantle phases, melts are relatively enriched in LREE while residues are depleted. Extraction of continental crust depleted residual mantle in incompatible elements continental crust residual after extraction from whole mantle residual after extraction from upper mantle Increasing crustal concentrations Increasing compatibility Crust data: Taylor and McLennan (1985); Primitive mantle (BSE) data: Hofmann (1988) Trace element spectra of basalts from different tectonic contexts Ocean basins Arcs White (1997) LIL LIL REE HFSE HFSE Increasing compatibility during peridotite melting REE Basalts from different tectonic contexts have different trace element spectra that reflect source composition, source phases, pressure, mechanism and extent of melting. Isotopic variations in the crust and mantle Radiogenic isotopes • Nuclear processes • Absolute dating • Powerful source tracers • Not fractionated by geochemical processes (or minor fractionation is corrected during analysis) Stable isotope fractionation • Physicochemical processes • Varies in a systematic manner with isotope mass • Sensitive to geochemical processes and temperature • Source tracers "Mass independant" isotope fractionation • Sensitive to highly specific processes Mechanisms of Radioactive decay 1) Emission of a β particle (Rb-Sr, Re-Os) β- Electron emission from nucleus. The atomic number increases by one, the mass remains unchanged. β+ Positron emission from nucleus. The atomic number decreases by one, the mass remains unchanged. 2) Electron capture (K-Ar) Nucleus absorbs an inner shell electron. The atomic number decreases by one, the mass remains unchanged. 3) Emission of an α particle (nucleus of He) (Sm-Nd) The atomic number decreases by two, the atomic mass decreases by four. 4) Spontaneous fission (U-Xe) 5) Decay chains (U-Pb, Th-Pb) The radiogenic daughters are themselves radioactive. A whole series of instable isotopes exists between the radioactive parent (238U, 235U or 232Th) and the stable daughter (206Pb, 207Pb, 208Pb). Nucleii of 4He atoms are produced by emission of α particles. Radiometric isotope systems • Systems with long half lives - used for dating and tracing sources of mantle and crustal magmas 87Rb 87Sr 147Sm 143Nd 176Lu 176Hf 187Re 187Os 40K 40Ar • U-Th-Pb decay series - provide time constraints for magmatic processes 238U ... 206Pb 235U ... 207Pb 232Th ... 208Pb • t1/2 = 4.47 Gy t1/2 = 0.7 Gy t1/2 = 14 Gy by-product 4He Cosmogenic isotopes 10Be • t1/2 = 48 Gy t1/2 = 106 Gy t1/2 = 36 Gy t1/2 = 42 Gy t1/2 = 1.28 Gy (from spallation) t1/2 = 1.4 My Short-lived extinct isotopes - provide constraints on earliest Earth processes and on material from which Earth formed 146Sm 142Nd 182Hf 182W t1/2 = 103 My t1/2 = 9 My Black: lithophile; Red: siderophile/chalcophile; Blue: volatile Radioactive decay equations The rate at which the abundance of a radioactive isotope (parent) decreases is proportional to the number of atoms remaining (N) : -dN/dt = λN λ = decay constant Solution: N = Noe-λt No = number of atoms of the parent at t = 0. D* = No- N = N(eλt - 1) D* = number of atoms of the daughter produced by radioactive decay D = Do + D* = Do + N(eλt - 1) Do = number of atoms of the daughter t = 0 D = total number of atoms of the daughter Radioactive decay equations (cont.) Example of Rb-Sr couple : 87Rb (parent) 87Sr (daughter) (t1/2 ~ 49 Gy) D = Do + N(eλt - 1) 87Sr = 87Sro + 87Rb(eλt - 1) Normalized to a non-radiogenic isotope (here 86Sr) because: • not all phases will have same 87Sro, but all should have same 87Sr/86Sr • it is much easier to precisely determine an isotopic ratio than the absolute quantity of a given isotope 87Sr/86Sr = (87Sr/86Sr)o + (87Rb/86Sr)(eλt - 1) Linear equation y-intercept (initial ratio) Slope (depends on age) Isochron method Example of Rb-Sr system 87Sr/86Sr t2 t1 sample 1 sample 2 t0 sample 3 87Rb/86Sr Evolution of 87Sr/86Sr over time to 86Sr, a non-radiogenic isotope.) (87Sr and 87Rb are normalized Conditions required to produce an isochron 87Sr/86Sr t2 t1 t0 87Rb/86Sr 1) At t0, all phases must have the same isotopic ratio 2) The range of parent/daughter ratios must be substantial 3) The system must remain closed since its formation 87Sr/86Sr = 87Sr/86Sri + 87Rb/86Sr (eλt - 1) 87Sr/86Sr t = age λ= decay constant of 87Rb = 1.42 x 10-11 years-1 87Sr/86Sr = initial isotopic ratio i Initial isotopic ratio = 87Sr/86Sr i Provides information about the source Isochron equation Slope = eλt - 1 The age is determined from the slope of the isochron 87Rb/86Sr How do radiogenic isotope compositions of major Earth reservoirs evolve through time? continental crust residual after extraction from whole mantle residual after extraction from upper mantle Increasing crustal concentrations Rb, Sr, Sm, Nd, Lu, H are all incompatible during mantle melting. But: DRb < DSr (Rb more incompatible than Sr) D= bulk partition coefficient DSm > DNd DLu > DHf Sr isotopic evolution of Continental Crust and Mantle (?) Since DRb< DSr Rb/Sr crust > Rb/Sr BSE > Rb/Sr residual mantle White (1997) BABI (Basaltic Achondrite Best Initial) = estimate of solar system initial 87Sr/86Sr Over time, the crust has developed a radiogenic signature relative to BSE while the residual mantle has developed an unradiogenic signature. Exact 87Sr/86Sr values of each reservoir depend on crustal extraction history. Nd isotopic evolution of Crust and Mantle Since DSm> DNd Sm/Nd crust < Sm/Nd BSE < Sm/Nd mantle (?) chondritic Sm/Nd As a result, residual mantle Nd becomes radiogenic while crustal Nd becomes relatively unradiogenic over time. " Nd ! $ 143 Nd ' sample ) ) & 144 ( Nd = 10000 * & 143 #1) Nd & chondrite) ) ( 144 % Nd ( CHUR = chondritic uniform reservoir What fraction of the mantle was depleted by formation of the Continental Crust? Can be calculated from the mean Nd isotopic compositions and concentrations of the continental crust, the depleted mantle (known from MORB) and the Bulk Silicate Earth. But our knowledge of each of these parameters, particularly for the BSE, is limited. (We'll come back to this...) (?) εNd and 87Sr/86Sr are anti-correlated among main Earth reservoirs White (1997) The opposing behavior of Nd and Sr isotopes is due to the fact the parent is more incompatible than the daughter in one case (Rb/Sr) and less incompatible in the other (Sm/Nd) Nd and Hf systematics in oceanic basalts Continental Crust Dickin (2000) In both the Sm/Nd and Lu/Hf systems the parent isotope is less incompatible than the daughter. Produces with time a positive correlation. Calculating Nd model ages for extraction of a crustal sample from the mantle Geologic significance of model age depends on validity of the model. Based on two questionable assumptions: 1) The mantle evolution curve is well defined. 2) The Sm/Nd ratio of the sample has remained unchanged since time of extraction from the mantle. Nevertheless, can provide approximate indication of age of crustal extracton. Calculating Model ages At the time of crustal extraction from the mantle, the sample will have the same Nd isotopic ratio (143Nd/144Ndi) as the mantle evolution curve at that time (here assumed to be the depleted mantle curve, DM). So: crustal sample depleted mantle " 143 Nd % " 143 Nd % " 147 Sm % (T $ 144 ' = $ 144 ' + $ 144 ' (e M )1) # Nd & DM # Nd & i # Nd & DM and " 143 Nd % " 143 Nd % " 147 Sm % (T $ 144 ' = $ 144 ' + $ 144 ' (e M )1) # Nd & DM # Nd & i # Nd & DM Combining equations: ! ! " 143 % " 143 Nd % " 147 Sm % " 147 Sm % Nd )T DM ( $ 144 ' (e (1) = $ 144 ' ( $ 144 ' (e )TDM (1) $ 144 ' # Nd & sample # Nd & sample # Nd & DM # Nd & DM After a little algebra: ! TDM " " 143 Nd % % " 143 Nd % $ $ 144 ' ( $ 144 ' ' Nd Nd # & # & $ ' DM sample ln$ 147 +1' 147 " Sm % " Sm % $$ $ 144 ' ( $ 144 ' '' Nd Nd # & # & # DM sample & = ) where λ = 6.54 x 10-12 yr-1 (decay constant of 147Sm) U-Th-Pb decay systems 235U, 238U and 232Th decay to Pb through a series of short-lived radioactive daughter isotopes. 235U ... 207Pb t1/2= 0.70 Gy 238U ... 206Pb t1/2= 4.47 Gy 232Th ... 208Pb t1/2= 14.0 Gy For long-term processes (e.g. evolution of mantle reservoirs), can be approximated as a direct decay from the parent (U or Th) to Pb. For relatively rapid processes (e.g. magma generation) intermediate daughters can be very useful for constraining rates 238U Decay chain • Several intermediate daughters with half-lives appropriate for studying magmatism (234U, 230Th, 226Ra) • Each α particle is a 4He nucleus Pb isotopes 204Pb, 206Pb, 207Pb, 208Pb 235U ... 207Pb t1/2= 0.70 Gy 238U ... 206Pb t1/2= 4.47 Gy Two isotope systems in which parent is an isotope of U and daughter is an isotope of Pb. Therefore can write an isochron equation which does not involve the U/Pb ratio. 232Th ... 208Pb t1/2= 14.0 Gy 204Pb, the only non-radiogenic Pb isotope, is used for normalization (206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb) Pb-Pb isochrons 207Pb/204Pb - (207Pb/204Pb)T = ( 235U/204Pb)(eλ235T - 1) 206Pb/204Pb - (206Pb/204Pb)T = ( 238U/204Pb)(eλ238T - 1) where T is age of event to be dated Dividing the two equations and replacing the ratio 235U/238U by its current value (1/137.88), uniform (or nearly so) in the solar system: 207Pb/204Pb - (207Pb/204Pb)T 206Pb/204Pb (206Pb/204Pb) - T = (eλ235T - l) 137.88 (eλ238T - l) Nevertheless, even if U doesn't appear in the equation, a substantial range of U/Pb is needed to generate an isochron. Example: Pb-Pb isochron of 3 Gy 16.0 t=2 Gy 15.8 t=1 Gy µ=10 15.6 207Pb/204Pb t=0 µ=9 15.4 µ=8 15.2 15.0 14.8 14.6 14.4 µ = 238U/204Pb 14.2 12 13 14 15 16 17 18 19 20 206Pb/204Pb To calculate Pb composition at a time t, following an event of age T: 207Pb/204Pb t - (207Pb/204Pb)T = (µ/137.88)∗(eλ235T - eλ235t) 206Pb/204Pb t - (206Pb/204Pb)T = µ∗(eλ238T - eλ238t) (µ/137.88 = 235U/204Pb) Holmes-Houtermans model (1946) 4 b 20 P µ= 207Pb/204Pb t - (207Pb/204Pb)T 206Pb/204Pb t - (206Pb/204Pb)T 8 / 23 U (eλ235T - eλ235t) = 137.88 (eλ238T - eλ238t) T= age of the Earth t = age of galena sample Determination of age of Earth with Pb-Pb isochron after Patterson (1956) Ocean sediments (which average large volumes of material) plot on isochron defined by meteorites. Lead Paradox White (1997) Pb isotopic compositions of most accessible Earth reservoirs, including MORB, plot to the enriched side of the Geochron. But average Bulk Silicate Earth is expected to plot on the Geochron. Where is the "missing" unradiogenic Pb? Core? Lower crust? Subducted slabs? Still not resolved. Another use of Pb isotopes: U-Pb dating of zircons Concordia diagram Extremely powerful tool for dating crustal rocks • Zircon: very common crustal phase containing almost no Pb but a lot of U • All Pb of radiogenic origin • Highly resistant phase • in situ techniques allow determination of different ages from different portions of zoned zircons In situ U-Pb dating of zircons bar = 50 µm numbers = age in My Fornelli et al. (2014) Allows different ages to be found for different portions of zircon grains (e.g., old xenocrystic cores and later metamorphic overgrowths). In situ U-Pb dating often accompanied by Hf isotopic measurements. In situ Hf isotopic measurements in zircon have become very common - Why? • Hf, like Zr, is a High Field Strength Element (HFSE), and so can substitute for Zr in zircon structure • As a result, Hf concentrations in zircon are very high (1% or more) and Lu/Hf ratios are very low. • Hf isotopes in zircon are excellent source tracers because low Lu/Hf ratios minimize radiogenic ingrowth of 176Hf after zircon crystallization • Hf isotope data can be coupled with U/Pb ages in zircon • Measurements by laser ablation ICPMS are rapid and "easy" (though several pitfalls). De p le 176Hf/177Hf BS E C ru s 0 te d ma n tl e Hf crustal model ages of zircons ("two stage" ages) t zircon crystallization age Hf Model age time (Gy) 4.5 stage 1: Evolution of Hf isotope composition of source rocks with an average crustal 176Lu/176Hf ratio (an assumed value often used) stage 2: Hf residence in zircon using measured 176Lu/176Hf ratio (~0) Major assumptions: • Crustal precursor from which zircon parental magma formed had an "average crustal" 176Lu/177Hf ratio. • Crust derived from a mantle with a well-defined "depleted" Hf isotopic signature. De 176Hf/177Hf BS p le E C ru s 0 te d ma n tl e Hf crustal model ages of zircons ("two stage" ages) t zircon Hf Model age crystallization age 4.5 D ep $ ' Hf & 177 ( sample) ) Hf = 10000 * & 176 #1) & Hf ) & 177 (chondrite) ) % Hf ( ! a n tl e BSE C ru s 176 " Hf dm εHf Same idea, but in εHf notation le te 0 t zircon Hf Model age crystallization age 4.5