Earth's mantle
Earth's mantle
Lower mantle (3)
Between the Earth's core and the lower mantle (depth: 2890 km, 136 GPa, 3500 °C) lies the core-mantle boundary (CMB) with the subsequent D″ layer (approximately 200–300 km thick). The D″ layer exhibits heterogeneous regions (detected by seismic measurements) known as Large Low Shear Velocity Provinces (LLSVPs) . These have a lower density than the Earth's core but a higher density than the lower mantle, thus influencing hot spots and mantle convection. They are explained by:
Reaction of mantle material with Earth's core material (mixed metal-silicate phase; Kito et al. 2004)
Subducted lithospheric plates sink into the layer (plate graveyard; Ding et al. 1997)
Transformation of a mineral phase into a different crystal structure with higher density (bridgmanite → post-perovskite; Oganov & Ono 2004)
The model composition of the entire mantle is assumed to be pyrolite (Ringwood 1962), with a surface component of 60% peridotite and 40% basalt/eclogite. Due to increasing pressure and temperature, phase transformations of the model rock occur with increasing depth. In particular, monoclinic/trigonal/tetragonal unit cells are transformed into orthorhombic ones. The larger unit cell allows for more efficient stacking and thus a lower packing density.
The more precise mineral composition of the Earth's mantle was detected through (1) high-pressure diamond anvil cell (LHDAC) experiments , (2) inclusions in sublithospheric diamonds brought to the surface by cratonic kimberlites, and (3) mineral formations in impact craters and meteorites.
The following main phases are present in the lower mantle (Wicks & Duffy 2016):
Bridgmanite : (Mg,Fe) SiO3 , approximately 80%, orthorhombic distorted perovskite structure, with corner-sharing SiO6 octahedra and Mg,Fe cations in the octahedral interstices (12-8 coordination); found in diamonds and meteorites
Ferropericlase : (Mg,Fe)O, mixed crystal of magnesiowüstite (MgO) and wüstite (FeO), proportion approx. 15%, face-centered cubic NaCl-type lattice with octahedral coordination for oxygen and Mg/Fe; found in diamonds
Ca-perovskite (davemaoite) : CaSiO3 , proportion approx. 10%, cubic perovskite structure with SiO6 octahedra and Ca cations in the octahedral voids (12 coordinates)
In addition, other secondary phases occur in smaller proportions, some of which were induced by subduction into the mantle (Kaminsky 2017, Wicks & Duffy 2016):
Post-perovskite : (Mg,Fe) SiO3 , only at the D″ layer (from bridgmanite, from approx. 125 GPa = 2600 km), orthorhombic CaIrO3- type structure with SiO6 octahedra in layers and cations in between (volume reduction to bridgmanite by 1-1.5%); found in high-pressure diamond anvil cell (HHDAC) experiments.
Stishovite : SiO₂ , stable from ~7–50 GPa, tetragonal rutile-type structure with corner-sharing SiO₆ octahedra; found in impact craters and meteorites
Post-stishovite : SiO₂ , stable from ~50–70 GPa, orthorhombic CaCl₂-type structure, with edge-sharing SiO₆ octahedra; found in impact craters and meteorites
Seifertite : SiO₂ , stable at > 70 GPa, orthorhombic α- PbO₂ -type structure, with edge- and corner-sharing SiO₆ octahedra (highest packing density); found in impact craters and meteorites, possibly only present at the D″ layer
Magnesite II (MgCO 3 ) which is first converted into an orthorhombic pyroxene structure and then into a perovskite structure (thus stable up to 200 GPa, beyond that → MgO + Diamond)
Post-aragonite ( CaCO3 ) with orthorhombic pyroxene structure (> 137 GPa)
Dolomite II and Dolomite III with as yet unknown structure (40 GPa)
Diamond : The main carriers of lower mantle inclusions are so-called "superdeep" or "sublithospheric" diamonds. These are formed by deep plate subduction, the release of COH fluids, or the reaction of fluids with carbonates from the lower mantle.
Other theoretically postulated phases include the NAL phase (sodium aluminum silicate: NaAlSi2O6 ), the CAS phase (calcium aluminum silicate: CaAl4Si2O11 ) , the CF phase (Ca ferrite: CaFe2O4 ) , and the PhH phase ( MgSiH2O4 ) .
Spin transition zone (STZ) : Transition of Fe²⁺ in bridgmanite and ferropericlase from a high-spin (HS) to a low-spin (LS) state, starting at pressures of approximately 50 GPa = 1200 km depth (Kaminsky 2017). Iron (Fe) is a transition metal with unpaired electrons in its d orbitals. In bridgmanite/ferropericlase, the iron is coordinated octahedrally to oxygen. The five d orbitals of Fe are energetically split into three lower-energy orbitals (t ≤ g) and two higher-energy orbitals (e ≤ g), a phenomenon known as crystal field splitting. The increasing pressure causes the distance between Fe and O to decrease, which strengthens the electric field of the O²⁻ ions on the Fe ion, resulting in a greater crystal field splitting. This results in more paired electrons being found in lower-energy t 2 g orbitals, fewer unpaired electrons, and thus a lower total spin. The electron configuration becomes more compact, and consequently, so do the ionic radii, leading to a higher mantle density. This density anomaly triggered by the transition is evident in seismic measurements and has direct effects on differentiation and mantle convection processes (Kaminsky 2017).
The lower mantle is the source of approximately 10–20% of hot spots (e.g., Hawaii). Geological markers (Courtillot et al. 2003) include, for example, (1) high 3He / 4He ratios (low differentiation), (2) seismic anomalies (LLSVPs), (3) high local stability of the hot spots, (4) higher temperatures, (5) lower 182W / 184W ratios indicating core-mantle interaction, (6) high Fe/Mn ratios, typical of a reducing environment, and (7) low 187Os / 188Os ratios.
The lower mantle (LM) has a thickness of 2,230 km (660 – 2,890 km), density: 4.5 – 5.6 g/cm 3 , temperature: 1,900 – 4,000 °C, pressure: 23 – 136 GPa (Dziewonski & Anderson 1981, Katsura 2025).
Transition zone
The mantle transition zone (MTZ) extends from 410 to 660 km depth (Frost 2008) and contains a number of seismic discontinuities (seismic velocity anomalies) resulting from mineral phase transformations (Agee 1998). The most significant transformation occurs at the boundary between the lower mantle and the mantle transition zone, the 660 km discontinuity (Zhang et al. 2021).
Bridgmanite (Mg,Fe)SiO 3 + ferropericlase (Mg,Fe)O ⇒ ringwoodite β(Mg,Fe) 2 [SiO 4 ]
Ringwoodite has a cubic spinel-type structure with the oxygen anions in cubic close-packed arrangements and the cations (Mg, Fe, Si) in tetrahedral and octahedral interstices (Price & Parker 1984). Ringwoodite is stable from 18–23 GPa (520–660 km). The next transformation therefore occurs at the 520-km discontinuity (Zhang et al. 2021).
Ringwoodite β(Mg,Fe) 2 [SiO 4 ] ⇒ Wadsleyite γ(Mg,Fe) 2 [SiO 4 ]
Wadsleyite is stable from 13 to 18 GPa (410 to 520 km) and has a modified orthorhombic spinel structure with Mg and Fe at three different octahedral lattice positions (M1, M2, and M3), Si at one tetrahedral position, and O at four lattice positions. One of the O atoms is not silicon-bound and can be hydrated (forming OH), creating a vacancy at the M3 cation position (2H + = Mg2 + ) to balance the charge (Das et al. 2018). At given pressures (> 9 GPa) and a water content > 0.25 wt%, the orthorhombic structure transforms into a monoclinic crystal structure through distortion (Wang et al. 2023).
The “nominally anhydrous minerals (NAMS)” can absorb up to 1.4 wt% water under mantle transition conditions as “Hydrous-Ringwoodite (h-Rw)” and “Hydrous-Wadsleyite (h-Wd)” (Pearson et al. 2014 and Fei & Katsura 2021, respectively) ⇒ Transition zone: reservoir for large quantities of water originating from subduction zones, because older (> 100 Ma), colder lithospheric plates (cold subduction slabs) bring hydrous phases (3-18 wt%), so-called “Dense Hydrous Magnesium Silicates (DHMS)”, into the mantle transition zone, perhaps even as far as the 660 km discontinuity and into the lower manel (PhH, Walter et al. 2015). The following were detected in high-pressure experiments (Ohtani et al. 2001, Xu et al. 2021, Ishii & Ohtani 2025):
Hydrous Phase A (PhA): Mg 7 Si 2 O 14 H 6
Super “Hydrous Phase B (sPhB, SUB): Mg 10 Si 3 O 18 H 4
Hydrous Phase D (PhD) = Phase G: MgSi 2 O 6 H 2
Hydrous Phase E (PhE): MgSiO 6 H 4
Hydrous Phase H (PhH): MgSiO 4 H 2
Subducting plates form Wadati-Benioff zones (increasingly deep hypocenters of earthquakes along a subducting plate). The WBZ extends to a depth of 700 km (Frohlich 1989). In the transition zone at approximately 500 km depth, there is an increase in these earthquakes (Omori et al. 2004), which are classified as “deep earthquakes” and are presumably generated by phase transitions of water-containing deep geological formations (DHMS): the “Extended Dehydration-induced Earthquake (EDIE)” hypothesis (Omori et al. 2004).
The second most abundant phase (~40%) after ringwoodite/wadsleyite (60%) in the transition zone is the garnet representative majorite Mg₃ (SiMg)( SiO₄ ). Majorite forms progressively from subducted pyroxenes (e.g., enstatite MgSiO₃ ) at depths of 300 to 500 km (10–17 GPa; Yoshino et al. 2008). Majorite itself decays from about 25 GPa into bridgmanite, Ca-perovskite, and NAl/CF phases (Stagno et al. 2023, Wijbrans et al. 2016).
In the mantle-transition zone, the stable quartz phase is stishovite ( SiO2 ).
The MTZ has a thickness of 250 km (410 – 660 km), density: 3.9 – 4.3 g/cm 3 , temperature: 1,400 – 1,900 °C, pressure: 13 – 23 GPa (Dziewonski & Anderson 1981).
literature
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