Incompatible element
In geochemistry, incompatible elements are trace elements characterized by very low partition coefficients (D ≪ 1) between common rock-forming minerals and coexisting silicate melts, causing them to preferentially partition into the liquid phase rather than being incorporated into crystallizing solids during magmatic processes such as partial melting and fractional crystallization.[1] This behavior results from their ionic properties—typically large radii and/or high charges—that make them poor fits for the crystal lattices of major mantle and crustal minerals like olivine, pyroxene, and plagioclase.[2] These elements are broadly classified into groups such as large-ion lithophile elements (LILE), including potassium (K), rubidium (Rb), cesium (Cs), strontium (Sr), and barium (Ba), as well as high field strength elements (HFSE) like niobium (Nb), tantalum (Ta), zirconium (Zr), and hafnium (Hf), and the light rare earth elements (LREE) such as lanthanum (La) and cerium (Ce).[1] During partial melting of the mantle, incompatible elements become highly enriched in the generated melt, with concentrations inversely proportional to the degree of melting (e.g., small melt fractions of 1-2% can enrich elements like La by factors of up to 50 relative to the source).[1] In fractional crystallization, they remain in the evolving residual liquid, further concentrating as solids are removed, which contributes to the overall enrichment of the continental crust in these elements over Earth's history.[1] Incompatible elements serve as critical tracers in petrology and geodynamics, enabling scientists to infer mantle source compositions, melting conditions, and tectonic settings through ratios like La/Sm or Nb/Zr that are relatively insensitive to later modifications.[1] For instance, mid-ocean ridge basalts (MORB) often display depletions in HFSE due to prior melt extraction in the mantle, while ocean island basalts (OIB) show enrichments reflecting deeper, less depleted sources.[1] Their study has profound implications for understanding planetary differentiation, as repeated cycles of melting and solidification have led to the progressive incompatible element depletion of the mantle and enrichment of the crust.[1]Fundamentals
Definition
In geochemistry, an incompatible element refers to a trace element that is not readily incorporated into the crystal lattices of major rock-forming minerals during igneous differentiation processes, owing to significant mismatches in ionic radius, valence, or coordination requirements relative to the available lattice sites.[2][3] This reluctance to substitute into solid phases results in the element's strong partitioning preference for the coexisting melt.[1] Key characteristics of incompatible elements include their low affinity for common silicate minerals such as olivine, pyroxene, and plagioclase, which dominate mantle and crustal lithologies.[4] Consequently, these elements become progressively enriched in the residual liquid as crystallization proceeds or during partial melting of source rocks.[5] This behavior is quantified by the mineral-melt partition coefficient, typically much less than unity for such elements.[1] The concept of incompatible elements emerged in mid-20th century petrology to characterize trace elements systematically excluded from early-crystallizing phases in evolving magmas, building on foundational principles of ionic substitution established earlier. Incompatibility remains a phase-specific property, contingent on the mineral assemblage and thermodynamic conditions, rather than an intrinsic feature of the element's overall abundance in a rock's bulk composition.[1]Partition coefficient
The partition coefficient, denoted as D, quantifies the distribution of a trace element between a solid phase and its coexisting liquid phase in geochemical systems, defined as D = \frac{C_s}{C_l}, where C_s is the concentration of the element in the solid and C_l is the concentration in the liquid.[1] This ratio assumes equilibrium partitioning and is fundamental for assessing how trace elements are incorporated into minerals relative to melts.[6] Interpretation of D values provides insight into element compatibility: values much less than 1 (typically D < 0.1) indicate incompatibility, meaning the element prefers the liquid phase; D \approx 1 suggests moderate compatibility; and D \gg 1 denotes high compatibility, with the element favoring the solid phase.[1] Several factors influence D, including ionic radius governed by Goldschmidt's rules, which predict substitution based on size and charge similarity to major cations in the crystal lattice; charge balance requirements; crystal field stabilization effects that alter energies for transition metals; and environmental conditions such as temperature and pressure.[7][1] Experimental determination of D involves synthesizing or analyzing natural mineral-melt pairs under controlled conditions, followed by measurement using techniques like electron microprobe analysis (EMPA) for major and minor elements or laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) for trace levels in both natural and experimental systems.[8] These methods ensure precise concentration ratios by targeting spots on crystals and adjacent glasses.[9] For multi-mineral assemblages, the bulk distribution coefficient D_{\text{bulk}} represents a weighted average, calculated as D_{\text{bulk}} = \sum (X_i \cdot D_i), where X_i is the modal proportion of mineral i and D_i is its individual partition coefficient.[10] This aggregate value accounts for the overall partitioning behavior in rocks or magmas composed of multiple phases.[1]Classification
Types of incompatible elements
Incompatible elements are classified primarily by their degree of incompatibility, which is determined by bulk partition coefficients (D) relative to the melting phase, where D << 1 indicates strong partitioning into the melt. Highly incompatible elements, with D < 0.01, exhibit rapid enrichment in even small degrees of partial melting due to their minimal incorporation into solid phases. Moderately incompatible elements, characterized by D values between 0.01 and 0.1, show less pronounced enrichment but still preferentially enter the melt over the residue. Additionally, incompatible elements can be distinguished as fluid-mobile, which are readily transported by aqueous fluids during metasomatic processes, versus fluid-immobile types that remain more immobile in such environments.[1][2] Chemical subgroups of incompatible elements are defined based on ionic properties that govern their partitioning behavior. Large ion lithophile elements (LILE) possess large ionic radii and low charges, rendering them incompatible in most common mantle minerals due to poor lattice fit. High field strength elements (HFSE), in contrast, feature high charges and small ionic radii, leading to incompatibility arising from challenges in achieving charge balance within mineral structures. Rare earth elements (REE) form another key subgroup, with light REE typically displaying high incompatibility owing to their larger sizes compared to heavy REE, which may show moderate compatibility in certain phases.[1][2] Several factors influence the typing of incompatible elements, including valence state, which alters ionic radius and charge balance (e.g., variable oxidation states can shift compatibility); hydration potential, particularly for elements that form hydrated bonds and enhance fluid mobility; and compatibility in accessory minerals, such as zircon, which can sequester HFSE despite their overall incompatibility in major phases. These factors collectively determine how elements partition during igneous processes.[1][2] The classification of incompatible elements has evolved from early qualitative groupings in the 1960s and 1970s, which relied on ionic radius and charge to categorize lithophile behaviors, to modern quantitative schemes that employ multi-element diagrams for visualizing relative incompatibilities and fractionation patterns. Seminal contributions, such as Gast's introduction of LILE in 1972, laid the groundwork for subgroup distinctions, while subsequent advancements incorporated lattice strain models for precise D predictions.[2][1]Examples of incompatible elements
In geochemistry, incompatible elements are those that are strongly partitioned into the melt phase rather than the solid residue during magmatic processes, with common examples including the alkali metals potassium (K), rubidium (Rb), and cesium (Cs); the alkaline earth elements barium (Ba) and strontium (Sr); large ion lithophile elements (LILE) such as uranium (U) and thorium (Th); high field strength elements (HFSE) like niobium (Nb), tantalum (Ta), zirconium (Zr), and hafnium (Hf); and light rare earth elements (REE) including lanthanum (La), cerium (Ce), and neodymium (Nd).[2] The incompatibility of LILE arises primarily from their large ionic radii, which exceed those of common cations like calcium (Ca²⁺) and sodium (Na⁺), preventing efficient substitution into the octahedral sites of silicate minerals such as olivine and pyroxenes.[2] In contrast, HFSE exhibit high charge-to-radius ratios (Z/r > 2), resulting in strong electrostatic repulsion and difficulty in achieving local charge balance within the lattices of typical mantle minerals, leading to very low partition coefficients (D) in these phases.[2] Light REE share similar exclusion mechanisms due to their larger ionic sizes relative to heavier REE, favoring melt enrichment during partial melting.[2] Incompatibility can vary with mineralogy and conditions; for instance, Sr has a partition coefficient near zero in olivine (D_{Sr}^{ol/melt} ≈ 0) but is compatible in plagioclase (D_{Sr}^{plag/melt} ≈ 1.8), though plagioclase is uncommon in the deep mantle.[11] Similarly, Nb is highly incompatible in most mantle phases like olivine (D_{Nb}^{ol/melt} ≈ 0.004) and clinopyroxene (D_{Nb}^{cpx/melt} ≈ 0.004), but it can be accommodated in accessory minerals such as rutile, where D values exceed 10.[12] The following table summarizes approximate partition coefficients (D = concentration in mineral / concentration in melt) for selected incompatible elements in key mantle minerals and bulk spinel peridotite (modal composition ≈ 55% olivine, 25% orthopyroxene, 15% clinopyroxene, 5% spinel). Bulk D values are weighted averages derived from mineral-specific D and modal abundances, typically << 0.1, confirming overall incompatibility. Values are compiled from experimental data at upper mantle pressures (1–3 GPa) and basaltic melt compositions.[11][12]| Element | Group | D_{olivine} | D_{clinopyroxene} | D_{bulk peridotite} |
|---|---|---|---|---|
| Rb | LILE | < 0.001 | 0.01–0.02 | ≈ 0.005 |
| Sr | LILE | ≈ 0 | 0.04–0.07 | ≈ 0.01 |
| Ba | LILE | < 0.001 | 0.001–0.003 | ≈ 0.001 |
| U | LILE/HFSE | < 0.001 | < 0.001 | < 0.001 |
| Th | LILE/HFSE | < 0.001 | < 0.001 | < 0.001 |
| Nb | HFSE | ≈ 0.004 | 0.005–0.06 | ≈ 0.01 |
| Ta | HFSE | ≈ 0.004 | 0.005–0.06 | ≈ 0.01 |
| Zr | HFSE | 0.003–0.005 | 0.03–0.13 | ≈ 0.02 |
| Hf | HFSE | 0.003–0.005 | 0.03–0.13 | ≈ 0.02 |
| La | Light REE | ≈ 0.0003 | 0.02–0.04 | ≈ 0.007 |
| Ce | Light REE | 0.0002–0.0003 | 0.02–0.11 | ≈ 0.01 |
| Nd | Light REE | ≈ 0.001 | 0.06–0.20 | ≈ 0.02 |