Radiogenic Isotope Geology G 214 2005 Lecture 1

Radiogenic Isotope Geology G 214, 2005 Lecture 1: Introduction

Modelled evolution of Rb/Sr in mantle and crust

Modelled evolution of Sm/Nd in mantle and crust

Fundamentals • Unstable isotopes decay to other nuclides • The rate of decay is constant, and not affected by P, T, X… • Parent nuclide = radioactive nuclide that decays • Daughter nuclide(s) are the radiogenic atomic products

Table of the elements

Radiogenic Isotope Systems • The radiogenic isotope systems that are of interest in geology include the following • • K-Ar Rb-Sr Sm-Nd Re-Os U-Th-Pb Lu-Hf These system develop through radioactive decay of parent isotopes to daughter isotopes.

DECAY TYPES Beta Decay 40 -> 40 + b- + + Q 20 Ca • 19 K – Where b- is the beta particle, is the antineutrino and Q stands for the maximum decay energy. Alpha Emission 238 -> 234 + He 4 + Q 90 Th 2 • 92 U – Where 2 He 4 is the alpha particle and Q is the total alpha decay energy Branched Decay 176 -> 176 via negative beta decay 72 Hf • 71 Lu 176 -> 176 positron decay or electron capture. 70 Yb • 72 Hf

Why is this useful? Isotopic variations between rocks and minerals due to 1. Mass fractionation (as for stable isotopes): only effective for light isotopes: H, He, C, O, S 2. Daughters produced in varying proportions resulting from previous event of chemical fractionation • 40 K 40 Ar by radioactive decay • Basalt rhyolite by FX (a chemical fractionation process) • Rhyolite has more K than basalt • more 40 K more • 40 Ar/39 Ar 40 Ar over time in rhyolite than in basalt ratio will be different in each 3. Time: the longer 40 K 40 Ar decay takes place, the greater the difference between the basalt and rhyolite will be

The Decay Constant • Over time the amount of the daughter (radiogenic) isotope in a system increases and the amount of the parent (radioactive) isotope decreases as it decays away. If the rate of radioactive decay is known we can use the increase in the amount of radiogenic isotopes to measure time. • The rate of decay of a radioactive (parent) isotope is directly proportional to number of atoms of that isotope that are present in a system. • Mathematically, this is expressed as; – -d. N/dt = l. N, – where N = the number of parent atoms and l is the decay constant – The negative sign means that the rate decreases over time

The Radiogenic Decay Equation • The previous relationship is used to derive the radioactive decay equation which has the following form D = Do + N (elt – 1) • This equation is used extensively in radiogenic dating studies • In principle, D and N are measurable quantities, while D o is a constant whose value can be either assumed or calculated from data for cogenetic samples of the same age. • If these three variables are known the above equation can be solved for t to give an “age” for the rock or mineral in question.

The Rb-Sr system • Strontium has four naturally occurring isotopes all of which are stable – 38 Sr 88, 38 Sr 87, 38 Sr 86, 38 Sr 84 • Their isotopic abundances are approximately – 82. 53%, 7. 04%, 9. 87%, and 0. 56% • However the isotopic abundances of strontium isotopes varies because of the formation of radiogenic Sr 87 from the decay of naturally occurring Rb 87 • Therefore the precise isotopic composition of strontium in a rock or mineral depends on the age and Rb/Sr ratio of that rock or mineral.

The Rb-Sr System • If we are trying to date a rock using the Rb/Sr system then the basic decay equation we derived earlier has the form Sr 87 = Sr 87 i + Rb 87(elt – 1) • In practice, it is a lot easier to measure the ratio of isotopes in a sample of rock or a mineral, rather than their absolute abundances. Therefore we can divide the above equation through by the number of Sr 86 atoms which is constant because this isotope is stable and not produced by decay of a naturally occurring isotope of another element.

The Rb-Sr System • This gives us the equation – 87 Sr/86 Sr = (87 Sr/86 Sr)i + 87 Rb/86 Sr(elt – 1) • To solve this equation, the concentrations of Rb and Sr and the 87 Sr/86 Sr ratio must be measured. • The Sr isotope ratio is measured on a mass spectrometer whilst the concentrations of Rb and Sr are normally determined by XRF or ICPMS.

First a little bit about Rb and Sr · · Rb behaves like K micas and alkali feldspar Sr behaves like Ca plagioclase and apatite (but not clinopyroxene) Rock Type Rb ppm K ppm Sr ppm Ca ppm Ultrabasic Basaltic High Ca granite Low Ca granite Syenite Shale Sandstone Carbonate Deep sea clay 0. 2 30 110 170 110 140 60 3 10 1 465 440 100 200 300 20 610 2000 180 40 8, 300 25, 200 42, 000 48, 000 26, 600 10, 700 2, 900 25, 000 76, 000 25, 300 5, 100 18, 000 22, 100 39, 100 302, 300 312, 400 29, 000

The Isochron Technique • Requires 3 or more cogenetic samples with a range of Rb/Sr – 3 cogenetic rocks derived from a single source by partial melting, FX, etc. – 3 coexisting minerals with different K/Ca ratios in a single rock

What can we learn from this? 1. After each period of time, the 87 Sr producing a new line 87 Rb in each rock decays to 2. This line is still linear but is steeper than the previous line. 3. We can use this to tell us two important things • The age of the rock • The initial 87 Sr/86 Sr isotope ratio

Determining the Age of a Rock

Determining the Age of a Rock

Let’s look now at the initial ratio

The initial ratio • How do we know if a series of rocks are co-genetic? • For rocks to be co-genetic, implies that they are derived from the same parent material. • This parent material would have had a single isotope value, ie the initial isotope ratio 87 Sr/86 Sr • Therefore, all samples derived from the same parent magma should all have the same 87 Sr/86 Sr isotope ratio • If they don’t, it implies that they are derived from a different parent source.

Other Systems • U-Th-Pb are used to date both metamorphic and igneous systems • During partial melting and fractional crystallisation of magma, U and Th are concentrated in the liquid phase and become incorporated into the more silica-rich products. • Therefore igneous rocks of granitic composition are more enriched in U and Th than basaltic or ultramafic rocks • Consequently, the continental crust has more U and Th than the upper mantle.

Using U-Pb data for determining ages

U-Th-Pb in Rocks • Which minerals in rocks contain U, Th and Pb? – – Zircon Monazite Titanite Allanite

Growth zoning in zircon grain

What is the problem here? • How do we tie the growth and breakdown of accessory phases to metamorphic reactions? • Careful study of thin-sections to clarify what metamorphic reactions are taking place and how the growth and breakdown of zircon and other accessory phases are linked to these reactions

Let’s come back to what we’ve been looking at in this course • Barrovian Metamorphism of Pelites – Chlorite to Biotite Zone • phengite + chlorite -> biotite + phengite-poor muscovite + quart z + H 2 O – Garnet Zone • chlorite + muscovite -> garnet + biotite + quartz + H 2 O – Sillimanite Zone • staurolite + muscovite + quartz -> garnet + biotite + sillimanite + H 2 O • No accessory phases mentioned!

A Recent Study…. • A SHRIMP U–Pb and LA-ICP-MS trace element study of the petrogenesis of garnet–cordierite– orthoamphibole gneisses from the Central Zone of the Limpopo Belt, South Africa • Ian S. Buick, Jörg Hermann, Ian S. Williams, Roger L. Gibson and Daniela Rubatto

Objectives of the study Constraining the timing of high-grade metamorphism is commonly difficult because age determinations that involve U–Pb geochronology of accessory phases (monazite, zircon, titanite, xenotime) need to be linked to pressure–temperature–time (P–T–t) paths through the growth/resorption of major P–T sensitive rockforming minerals such as garnet. Making this link is non-trivial in high-grade rocks because accessory phases may: 1) pre-date and be unrelated to metamorphism (e. g. , detrital zircon); 2) form during prograde metamorphic reactions involving other accessory phases and/or silicate minerals; 3) form at the metamorphic peak; 4) precipitate after the metamorphic peak, during the crystallisation of partial melt or exsolution of fluid at the solidus; 5) form along the retrograde P–T–t path due to breakdown of major rock-forming minerals; or 6) form due to solid state recrystallisation of preexisting grains at any time during metamorphism.

What did they conclude? • Garnet shows little major element zoning but significant trace element zonation characterised by core to rim increases in Y and the HREE and decreases in Sc and the LREE, followed by minor increases within 100 μm of the rim due to incipient garnet resorption. Zoning patterns for Y and the HREE were controlled by the breakdown of accessory phases (allanite and xenotime ± apatite). • Allanite + Xenotime + Sillimanite = Monazite + Garnet + Quar tz + H 2 O • Used this information to accurately constrain the timing of high-grade metamorphism in this part of the Limpopo Belt.