Thermal Equations of State of Magnesite: Implication for the Complex Mid‐Lower Mantle Seismic Scatterers

Magnesite (MgCO3) entering the lower mantle together with the subducted oceanic crust is an important carbon carrier. The reaction between magnesite and mantle minerals has been documented, but its influence on the density and velocity profiles of lower mantle remains unexplored. To decipher the deep carbon transportation and its associated effect, here we determined the thermal equations of state of magnesite up to 120 GPa and 2600 K using X‐ray diffraction in laser‐heated diamond anvil cells. The obtained thermal elastic parameters of magnesite facilitated a comprehensive understanding on the influence of magnesite‐SiO2 reaction, variation of carbon and SiO2 content, and temperature on the origin of lower‐mantle scatterers at 1,000–1,800 km depth. Our modeling revealed that the depth of the lower‐mantle VS scatterers is mainly controlled by the Al2O3 content in SiO2, while its magnitude depends on the SiO2 content. Along normal geotherm, the magnesite‐SiO2 reaction would occur before the post‐stishovite transition, consuming substantial SiO2 in the subducted oceanic crust. Depending on the amount of residual SiO2, the post‐stishovite transition can produce a 2.5–5.2 (2)% VS reduction, compatible with the observed seismic scatterers in Izu‐Bonin and Mariana subduction zones. Along slab geotherm, this reaction occurs after the post‐stishovite transition, generating a greater VS reduction of 4.4–6.4 (4)%. We thus propose that the reaction between sinking MgCO3 and SiO2 in the slab is one of the potential factors influencing the magnitude of the lower‐Vs scatterers at 1,000–1,900 km depth. Our results provide new insights into the deep‐mantle carbonate transportation influencing regional geophysics.


Introduction
The subduction slab is a bridge connecting the Earth's surface and interior (Bekaert et al., 2021).A large amount of surface material, especially volatile components such as carbon and water, will enter the Earth's interior with the sinking of the subduction slab (Bekaert et al., 2021;Dasgupta & Hirschmann, 2010;Walter et al., 2011).During this descent, these carbon and water-containing material continuously degas due to the increasing temperature and pressure (Dasgupta & Hirschmann, 2010).This not only causes partial melting of the overlying mantle wedge, affects the migration and distribution of elements within the Earth's interior, but also reacts with materials in the Earth's interior, leading to lateral heterogeneity in mantle composition (Zheng, 2019).Therefore, studying the physical properties of materials entering the Earth's interior with subduction slab under high pressure-temperature conditions is of great significance for understanding the circulation of materials in the deep Earth and the mantle lateral composition heterogeneity.
Meanwhile, MgCO 3 was reported to react with SiO 2 or davemaoite (CaSiO 3 ) when it subducted into the midlower mantle (Drewitt et al., 2019;Lv et al., 2021;Maeda et al., 2017).For example, prior to transitioning into magnesite-II, magnesite reacts with SiO 2 to form bridgmanite and CO 2 between 40 and 80 GPa with a Clapeyron slope of 74 to 147 MPa/K (Drewitt et al., 2019;Maeda et al., 2017).Residue magnesite will transition into magnesite-II above 80 GPa and 1700 K and further react with SiO 2 to form bridgmanite and diamond (Maeda et al., 2017).A number of small-scale scatterers have been identified at depths of 900-1,900 km with 2% to 12% shear velocity reduction, which were previously proposed to be associated with the structural phase transition of SiO 2 (e.g., Haugland et al., 2017;Hirose et al., 2005;Kaneshima, 2016;Kaneshima, 2019;Li & Yuen, 2014).The reaction between magnesite and SiO 2 in the mid-lower mantle not only consumes magnesite from the subducting oceanic crust but also affects the content of SiO 2 transported from the subducting oceanic crust to the lower mantle.The impact of such reactions on the formation of small-scale seismic scatterers in the mid-lower mantle remains unclear, and further in-depth research on the thermal elastic parameters of magnesite up to lower mantle condition could lead to a deeper understanding of this issue.
Although the physical properties of magnesite have been studied using various experimental techniques, the thermoelastic parameters have only been derived either from experiments at pressures less than 32 GPa up to 2073 K or only under 300 K (Fiquet & Reynard, 1999;Litasov et al., 2008;Redfern et al., 1993;Ross, 1997;Stekiel et al., 2017;Zhang et al., 1997).The isothermal bulk modulus, K T0 , determined by previous X-ray diffraction (XRD) experiments at 300 K ranges between 108 and 151 GPa, which exhibits strong tradeoff with its pressure derivative K' of 2.3-4.6 (Litasov et al., 2008;Redfern et al., 1993;Ross, 1997;Zhang et al., 1997).More importantly, the influence of temperature on the density and elasticity of magnesite has been investigated only below 32 GPa (Litasov et al., 2008;Zhang et al., 1997).Extrapolating the density and sound velocity of magnesite determined at low pressures to the relevant pressure conditions of the deep lower mantle would result in large uncertainties.
In this study, we conduct high pressure and temperature (P-T ) XRD experiments on magnesite with laser-heated diamond anvil cells (DACs) up to 120 GPa and 2600 K. Using the fitted thermal EoS of magnesite, along with previous experimental data for mantle minerals, we have modeled the density (ρ), shear velocity (V S ), and compressional velocity (V P ) profiles of magnesite.Furthermore, we discussed the influence of the magnesite + SiO 2 reaction on the ρ, V S and V P structure of subducted oceanic crust under lower mantle P-T conditions.These modeled density and velocity profiles are used to understand the seismic signature of carbonates and the influence of decarbonation on the seismic profiles along subducting slabs at the depth of mid-lower mantle.

Experiments
The natural single-crystal magnesite used in this study was from the Vargas Mineral Collection at the Jackson School of Geosciences, the University of Texas at Austin (Collection number: V3782).The composition of the MgCO 3 sample was examined by electron microprobe analysis, with less than 0.5 mol.%Mn and Fe in total (Yang et al., 2014).XRD measurements under ambient conditions revealed a rhombohedral structure for the sample (space group: R-3c) with a volume of 280.71 Å 3 (Yang et al., 2014).The sample was ground by agate mortar into fine powder.5 wt.%Pt was mixed into the sample powder as the heat absorber and pressure calibrant (Fei et al., 2007).The magnesite-Pt mixture was pressed into ∼10 μm thick foils, and cut into pieces of 40-80 μm afterward.The sample pieces were then sandwiched by two pieces of NaCl or KCl of ∼5-7 μm in thickness as the pressure medium and thermal insulator, and Re was used as the gasket material.Experiments below 90 GPa were performed by using DACs equipped with a pair of 300-μm or 200-μm culet diamonds, while data collected above 90 GPa were obtained by using DACs with two beveled diamonds (150-μm inner culet).
The high P-T XRD experiments were performed at the GeoSoilEnviroConsortium (GSECARS) of the advanced photon source (APS), argonne national laboratory (ANL).The incident X-ray beam with a wavelength of 0.3344 Å was ∼3 × 4 μm in size.The sample was initially compressed to ∼30 GPa at room temperature in which XRD patterns were collected every 2-5 GPa step (Table S1).The sample was heated from both sides (upstream and downstream) using double-sided flat-top infrared lasers, from 30 GPa.The temperature of the heated sample was determined by fitting the collected thermal radiation spectrum with the Planck radiation function under the Graybody approximation (Prakapenka et al., 2008), while pressures of the heated sample were determined by the thermal EoS of Pt (Fei et al., 2007).The temperature uncertainty was ∼10% of the measured value by considering the difference in temperature between the upstream and downstream heating sides and the temperature gradient inside the X-ray spot.During each heating cycle between 30 and 120 GPa, XRD patterns were collected at every 200 K from 1200 to 2600 K. XRD patterns were also collected after temperature quenching at high pressure.Dioptas was used to process the XRD data initially and to determine the occurrence of the phase transition (Prescher & Prakapenka, 2015).Origin or Igor program was then used to fit the peak positions, which are used to determine the lattice parameters of magnesite.

Result
At temperatures below 2200 K between 33 and 120 GPa, all of our collected XRD patterns can be well interpreted by magnesite, pressure calibrant Pt, and pressure medium KCl (Figures 1, 2 and Figure S2 in Supporting Information S1).The diffraction pattern under 73 GPa and 2000 K was refined using GSAS to confirm the structure of the sample (Figure S2 in Supporting Information S1) (Toby, 2001).Upon heating the sample to 2600 K and 80 GPa, we observed the appearance of new peaks, accompanied by a decrease in the intensity of MgCO 3  2020); gray dash lines: decomposition lines of MgCO 3 (Fiquet et al., 2002); gray dash-dotted lines: liquidus of MgCO 3 (Solopova et al., 2015); gray solid lines: normal mantle geotherm and a representative cold slab geotherm (500 K lower) (Katsura et al., 2010); open symbol with half red and half blue: coexistence of magnesite and magnesite-II.diffraction peaks (Figure 2).As we continued to increase the heating power, a significant laser flash occurred, indicating a rapid transformation.After quenching, the sample was observed to consist of a mixture of magnesite and the newly formed phase.Upon increasing pressure to 120 GPa between 2600 and 3000 K, the XRD patterns were composed of appearance of the new phase together with pressure calibrant and medium (Figure 2).No decomposition or melting of MgCO 3 was observed during all the heating cycles.Calculated deviatoric stress of our sample using collected diffraction patterns of Pt is less than 0.8 GPa at the experimental P-T range (Figure S1 in Supporting Information S1) (Dorfman et al., 2012).
The Pressure-Volume-Temperature (P-V-T ) data up to 120 GPa between 300 and 2,200 K were used to constrain the thermal EoS of magnesite (Table S1 in Supporting Information S1, Figures 1 and 3).The P-V-T data was fitted using both third-order isothermal Birch-Murnaghan and Mie-Grüneisen EoS (Table 1).The third-order isothermal Birch-Murnaghan EoS was expressed as: The K T and V 0 would be derived as follows: where K T0 and V 0 is the bulk modulus and volume under ambient condition; α is the thermal expansion, and dK T / dT is the pressure derivative of the bulk modulus.
The Mie-Grüneisen EoS enabled a more reliable extrapolation of the density and velocity profiles beyond the range of the experiments (Jackson & Rigden, 1996).The pressure could be described as

Table 1
Thermal Elastic Parameters of Magnesite where P c is the pressure at a reference temperature (T 0 = 300 K) from the Birch-Murnaghan EoS (Equation 1), and Pth is the thermal pressure caused by increasing temperature.Pth is calculated following (Jackson & Rigden, 1996): where Eth is the internal thermal energy resulting from the temperature increasing from T 0 to T, and γ is the Grüneisen parameter.Eth could be expressed with the Debye temperature, θ D : where n is the number of atoms of the magnesite formula and R is the gas constant, and the Grüneisen parameter γ could be derived as: where γ 0 is the Grüneisen parameter at ambient conditions and q is fixed as a constant.θ D can be calculated as follow: The fitting residuals with both Birch-Murnaghan and Mie-Grüneisen EoS are shown in Figure S3 of Supporting Information S1 to show the fitting quality of the data.

Discussion
The phase transition of magnesite has been the subject of extensive research for years (e.g., Binck et al., 2020;Boulard et al., 2011;Isshiki et al., 2004;Maeda et al., 2017).In this study, we observed the presence of new XRD peaks at 80 GPa when heating magnesite at ∼2600 K, but we did not observe the complete transition from magnesite to magnesite-II, even at temperatures above 2900 K within the pressure range of 80-120 GPa.Previous studies have also reported a wide coexistence range of magnesite and magnesite-II, extending up to 3100 K, which could be attributed to the limited heating time and the high kinetic energy barrier of the phase transition (Binck et al., 2020;Isshiki et al., 2004;Maeda et al., 2017).Since the motivation of this study was to determine the thermal EoS of magnesite, our heating time here was also limited to less than 5 minutes.

Journal of Geophysical
To examine the combined effect of pressure and temperature on ρ and V Φ of magnesite, we modeled them along normal mantle geotherm and 500-K colder slab geotherm, respectively (Figure 5) (Katsura et al., 2010;Litasov et al., 2008;Zhang et al., 1997).Our calculation has revealed that elevating temperature by 500 K can result in a 1.1-1.9(3)% and 1.3-1.7 (3)% reduction in the ρ and V Φ at lower-mantle pressures, respectively.Our calculated ρ below 50 GPa, as well as the temperature influence on ρ, aligns well with the findings of Litasov et al. (2008) within the range of calculation uncertainties.However, our ρ demonstrates a stronger dependence on pressure and is 1.6 (2)% higher than that reported by Litasov et al. (2008) at 85 GPa.It is worth noting that the experimental data in Litasov et al. (2008) were obtained at pressures below 32 GPa.A limited pressure range typically results in a larger pressure derivative of the bulk modulus and a lower bulk modulus, which can account for the relatively low ρ at 85 GPa when using the thermal elastic parameters from Litasov et al. (2008).Furthermore, a larger pressure derivative of the bulk modulus and a lower bulk modulus can also result in a greater V Φ in Litasov
Meanwhile, the variation in the depth of the observed scatterers in the midlower mantle could be explained by the varying amounts of Al 2 O 3 present in the SiO 2 (Zhang et al., 2022).It has been demonstrated that the addition of 1 mol.% of Al 2 O 3 can reduce the transition pressure from stv to post-stv by approximately 30 GPa (∼750 km), while 3 mol% of Al 2 O 3 can advance it by around 52 GPa (∼1,200 km) along the mantle geotherm (Zhang et al., 2022).It should be noted that the estimated low V S anomaly caused by the SiO 2 phase transition in the mid-lower mantle is much larger than the observed 2%-6% values in certain regions at depths of 1,000-1,600 km, such as the Japan Sea and Mariana (Kaneshima & Helffrich, 1999;Niu, 2014).Here, we have conducted a detailed study to investigate the impact of the MgCO 3 -SiO 2 reaction, SiO 2 content, SiO 2 phase transition, and temperature on the density and velocity structure of the subducted oceanic crust.The subducted oceanic crust primarily consists of mid-ocean ridge basalt (MORB).For our modeling, we considered a representative mineral assemblage of MORB, which consists of 39 vol.% bridgmanite, 30 vol.% davemaoite, 16 vol.%SiO 2 and 15 vol.%CF-type phase (Hirose et al., 2005;Ishii et al., 2022;Ricolleau et al., 2010).Then, we add 0-5 wt.% CO 2 in the form of MgCO 3 (0-12.1 vol.%) into the MORB as the carbonated oceanic crust and assume a complete reaction between MgCO 3 and SiO 2 (Dasgupta & Hirschmann, 2010).This allows us to evaluate the maximum effect of this reaction on the velocity and density structure of the carbonated oceanic crust.The thermoelastic parameters of the minerals used in the calculations can be found in Table S2 in Supporting Information S1, and the details of the modeling were included in Text S1 of Supporting Information S1.Since the effect of Al content on the V S reduction across the transition from stv to post-stv is within 3%, and the influence of pressure and temperature on the elasticity of Albearing SiO 2 is lacking, we only considered the influence of Al 2 O 3 on the depth of SiO 2 phase transition (Zhang et al., 2022).
Along the normal mantle geotherm, ρ, V P , and V S of the normal subducted oceanic crust with 16 vol.%SiO 2 are 2.3 (3)%, 0.6 (2)%, and 0.3 (2)% greater than those of the pyrolitic mantle at ∼50 GPa (∼1,250 km depth), respectively (Figure 6) (Katsura et al., 2010).Elevating the SiO 2 content from 16 to 25 vol.% can increase the velocity of the subducted oceanic crust by ∼1%-2% but rarely change the density (Figures 6 and 7).The difference in sound velocities between the subducted oceanic crust and pyrolitic mantle, particularly for V S , decreases with increasing depth when SiO 2 approaches the phase transition.V S of the subducted oceanic crust with  et al. (1997); thick lines: along normal mantle geotherm; thin lines: along a slab geotherm 500 K colder than the normal mantle (Katsura et al., 2010).SiO 2 content of 25 vol.%becomes indistinguishable from that of the pyrolitic mantle at ∼60 GPa (∼1,400 km depth) considering the calculation uncertainties (Figure 7).The phase transition from stv to post-stv in the subducted oceanic crust produces a pronounced softening in the V S .When the SiO 2 content is 16 vol.%,the maximum difference in the V S between the oceanic crust and pyrolitic mantle is 4.6 (3)% across the SiO 2 phase transition.Increasing the SiO 2 content to 25 vol.% in the subducted oceanic crust can enhance the low V S anomaly to 6.3 (4)% across the SiO 2 phase transition.In contrast, this phase transition exhibits a minor impact on V P and a negligible effect on ρ (Figures 6 and 7).V P of oceanic crust is only 0.7%-1.2%lower than the pyrolitic mantle across the SiO 2 phase transition.

Journal of Geophysical
Along normal mantle geotherm, the addition of carbon can effectively lower the ρ and sound velocity of the oceanic crust (Katsura et al., 2010).At 16 vol.%SiO 2 content, the presence of 2.5 wt.% CO 2 can lower the ρ, V P and V S of oceanic crust by ∼1.5 (4)%, 0.7 (1)%, and 1.2 (3)% compared to the carbon-free oceanic crust at 35 GPa (∼950 km), respectively (Figure 6).Increasing the CO 2 content to 5 wt.% makes the ρ and V P slightly lower than the pyrolitic mantle, and the V S anomaly reaches up to 1.6 (2)% (Figure 6).The MgCO 3 -SiO 2 reaction along the normal mantle geotherm occurs at ∼40 GPa (∼1,000 km), which is shallower than the depth of the post-stishovite transition and will consume large amount of SiO 2 in the subducted oceanic crust (Figure 6) (Katsura et al., 2010; Figure 6.The ρ, V P , and V S profiles of subducted oceanic crust with 16 vol.%SiO 2 along the normal mantle geotherm.(a), (b) and (c) are the modeled ρ, V P , and V S ; (d), (e) and (f) are the ρ, V P , and V S contrasts between normal or carbonated oceanic crust and pyrolitic mantle [ΔM = (M OC M Pyrolite )/M Pyrolite × 100%] (OC: oceanic crust); orange solid lines: carbonated oceanic crust with 5.0 wt.% CO 2 ; orange dashed lines: carbonated oceanic crust with 2.5 wt.% CO 2 ; blue solid lines: normal oceanic crust; gray solid lines: PREM.Orange shades show the position of the reaction between MgCO 3 and SiO 2 , while blue shades show the pressure range of the phase transition of SiO 2 .Stv: stishovite; post-stv: post-stishovite.Geotherm: from Katsura et al. (2010).Maeda et al., 2017).As a combined effect, the V P and V S of the subducted oceanic crust become indistinguishable from the pyrolitic mantle at 40-60 GPa (∼1,000-1,450 km) considering calculation uncertainties, while the density is ∼2% larger.
The reduction in the SiO 2 content in the subducted oceanic crust due to the MgCO 3 -SiO 2 reaction significantly lowers the magnitude of the V S softening but has a minor effect on V P and ρ when stv transitions into post-stv at ∼80 GPa (∼1,800 km) (Katsura et al., 2010;Zhang et al., 2021).At an initial SiO 2 content of 16 vol.%and CO 2 content of 2.5 wt.%, decarbonated oceanic crust exhibits 0.5 (1)% and 3.5 (4)% lower V P and V S , respectively, compared to the pyrolitic mantle at 80 GPa (Figure 6).Elevating the initial CO 2 content to 5 wt.% leads to an increased SiO 2 consumption during the MgCO 3 -SiO 2 reaction.As a result, the low V S anomaly at 80 GPa is only 2.5 (3)% (Figure 6).With an increased initial SiO 2 content of 25 vol.%, the MgCO 3 -SiO 2 reaction leaves behind 21 to 16 vol.%SiO 2 in the subducted oceanic crust by varying the initial CO 2 content from 2.5 to 5 wt.%.Consequently, the low V P and V S anomaly resulting from the post-stishovite transition in the subducted oceanic crust is 0.5-0.8(1)% and 4.1-5.2(6)% at 80 GPa, respectively, depending on the initial CO 2 content (Figure 7).It  c) are the modeled ρ, V P , and V S ; (d), (e) and (f) are the ρ, V P , and V S contrasts between normal or carbonated oceanic crust and pyrolitic mantle [ΔM = (M OC M Pyrolite )/M Pyrolite × 100%] (OC: oceanic crust); Orange solid lines: carbonated oceanic crust with 5.0 wt.% CO 2 ; orange dashed lines: carbonated oceanic crust with 2.5 wt.% CO 2 ; blue solid lines: normal oceanic crust; gray solid lines: PREM.Orange shades show the position of the reaction between MgCO 3 and SiO 2 , while blue shades show the pressure range of the phase transition of SiO 2 .Stv: stishovite; post-stv: post-stishovite.Geotherm: from Katsura et al. (2010).is noteworthy that the depth of the post-stishovite transition is strongly influenced by the Al 2 O 3 content in SiO 2 (Zhang et al., 2022).Variations in Al 2 O 3 content within SiO 2 from 5.2 wt.% to 0 can shift the previously mentioned low V S anomaly from 1,000 to 1,800 km depth (Figure 9).As a consequence, the combined effect of the MgCO 3 -SiO 2 reaction and Al 2 O 3 content in SiO 2 can provide an explanation for the observed 2%-6% low V S scatterers with a small V P anomaly in regions such as Japan Sea, Izu-Bonin, and Mariana within the mid-lower mantle between 1,000 and 1,600 km depth (Kaneshima & Helffrich, 1999;Niu, 2014;Niu et al., 2003).Other factors need to be considered to explain up to 9% density anomalies observed in the Mariana subduction zone (Niu et al., 2003).

Journal of Geophysical
Meanwhile, the depth of the MgCO 3 -SiO 2 reaction strongly depends on the slab temperature (Drewitt et al., 2019;Maeda et al., 2017).When the temperature of the subducted oceanic crust is ∼500 K lower than the surrounding mantle, the MgCO 3 -SiO 2 reaction will occur at ∼74 GPa, which is below the depth of the post-stishovite transition (Drewitt et al., 2019;Maeda et al., 2017) (Figure 8 and Figure S4 in Supporting Information S1).Since the addition of CO 2 will offset the density and sound velocities of the subducted oceanic crust to lower values, Figure 8.The ρ, V P , and V S profiles of subducted oceanic crust with 16 vol.%SiO 2 along a 500-K colder geotherm.(a), (b) and (c) are the modeled ρ, V P , and V S ; (d), (e) and (f) are the V P , and V S contrasts between normal or carbonated oceanic crust and pyrolitic mantle [ΔM = (M OC M Pyrolite )/M Pyrolite × 100%] (OC: oceanic crust); orange solid lines: carbonated oceanic crust with 5.0 wt.% CO 2 ; orange dashed lines: carbonated oceanic crust with 2.5 wt.% CO 2 ; blue solid lines: normal oceanic crust; gray solid lines: PREM.Orange shades show the position of the reaction between MgCO 3 and SiO 2 , while blue shades show the pressure range of the phase transition of SiO 2 .Stv: stishovite; post-stv: post-stishovite.Geotherm: from Katsura et al. (2010).carbonated oceanic crust with 16-25 vol% SiO 2 and 2.5% CO 2 will have the V S 4.4-5.9(7)% lower than the pyrolitic mantle.For a greater initial CO 2 content of 5 wt.%, the low V S anomaly produced by the post-stishovite transition in the subducted oceanic crust will be as large as 4.9-6.4(7)% (Figure 8 and Figure S4 in Supporting Information S1).
Our model here provides a detailed analysis of the impact of diverse factors, including carbon content, SiO 2 content, SiO 2 structural phase transitions, and temperature fluctuations within the subducted oceanic crust, on the density and velocity structure of the mid-lower mantle (Figure 9 and Figure S5 in Supporting Information S1).The variations in these factors provide an explanation for the observed small-scale scatterers, featuring varying magnitudes of low V S anomalies within the mid-lower mantle.We recognized that the low V S anomalies greater than 7% detected in the regions such as northeast China and South America at depths of 950-1,750 km in the lower mantle cannot be explained by our models (Haugland et al., 2017;Zhang et al., 2020).Since experimental constraints on the influence of Al 2 O 3 on the density and velocity of SiO 2 are lacking, whether the formation of these mid-lower mantle low-velocity scatterers is associated with a variation in the Al 2 O 3 content within SiO 2 in the subducted oceanic crust or caused by other factors requires further investigation in future studies.

Conclusion
In this study, we have determined thermal EoS of magnesite at relevant P-T conditions of the Earth's lower mantle.The phase transition from magnesite to magnesite-II was observed between 80 and 120 GPa above 2600 K, which is consistent with previous studies (Binck et al., 2020;Isshiki et al., 2004;Maeda et al., 2017).Magnesite in the subduction oceanic crust would react with SiO 2 and significantly affect the low V S anomaly caused by the phase transition of SiO 2 from stv to post-stv.Along a normal mantle geotherm, the magnesite-SiO 2 reaction will occur above the depth of the SiO 2 phase transition.As a result, varying the SiO 2 and CO 2 content in the subducted oceanic crust can lead to a change in the magnitude of the low-velocity anomaly caused by the SiO 2 phase transition from 2.5 to 5.2 (2)%.In contrast, the magnesite-SiO 2 reaction will occur below the depth of the SiO 2 phase transition along a slab geotherm.In this case, the V S anomaly caused by the SiO 2 phase transition will  (Zhang et al., 2021); blue contour: the velocity anomaly due to the post-stishovite transition in the subducted oceanic crust with a variation in the initial CO 2 content from 2.5 to 5.0 wt.%; gray dashed line: depth of the MgCO 3 -SiO 2 reaction along normal mantle geotherm (Drewitt et al., 2019;Maeda et al., 2017).Color vertical bars are the observed seismic low V S anomalies in the mid-lower mantle.Purple: northeast China (Zhang et al., 2020); yellow: Japan Sea (Niu, 2014); orange: Izu-Bonin (Niu et al., 2003); blue: Mariana (Kaneshima & Helffrich, 1999); green: South America (Haugland et al., 2017).
be less affected by the magnesite-SiO 2 reaction and would be 4.4-6.4(4)% depending on the initial SiO 2 and CO 2 content in the subducted oceanic crust.Considering a vary in the Al 2 O 3 content on the phase transition depth of SiO 2 , the aforementioned low-velocity anomaly can be present between the depth of 1,000 and 1,800 km in the lower mantle, which provided an explanation of low V S scatterers in various regions of the lower mantle, such as Japan Sea, Izu-Bonin and Mariana (Kaneshima & Helffrich, 1999;Niu, 2014;Niu et al., 2003).

Figure 2 .
Figure 2. Representative XRD patterns of magnesite at high pressures and temperatures.Black line: diffraction patterns.Vertical ticks indicate the diffraction peak of different material.Blue: Pt; orange: KCl; red: magnesite; purple: Re; green: new peaks appeared after phase transition.Incident X-ray wavelength is 0.3344 Å.

Figure 9 .
Figure 9. Low V P and V S anomalies generated by the carbonated oceanic crust across the post-stishovite phase transition along the normal mantle geotherm.(a) Carbonated oceanic crust with 16 vol.%SiO 2 ; (b) Carbonated oceanic crust with 25 vol.%SiO 2 .Gray arrow: depth of the post-stishovite transition due to the variation of Al 2 O 3 content in SiO 2(Zhang et al., 2021); blue contour: the velocity anomaly due to the post-stishovite transition in the subducted oceanic crust with a variation in the initial CO 2 content from 2.5 to 5.0 wt.%; gray dashed line: depth of the MgCO 3 -SiO 2 reaction along normal mantle geotherm(Drewitt et al., 2019;Maeda et al., 2017).Color vertical bars are the observed seismic low V S anomalies in the mid-lower mantle.Purple: northeast China(Zhang et al., 2020); yellow: Japan Sea(Niu, 2014); orange: Izu-Bonin(Niu et al., 2003); blue: Mariana(Kaneshima & Helffrich, 1999); green: South America(Haugland et al., 2017).