The termination of Neoproterozoic “Snowball Earth” glaciations is marked globally by laterally extensive neritic cap carbonates directly overlying glacial diamictites. The formation of these unique deposits on deglaciation calls for anomalously high calcium carbonate saturation. A popular mechanism to account for the source of requisite ocean alkalinity is the shallow-ridge hypothesis, in which initial spreading ridges surrounding fragments of Rodinia, assumed to be dominated by volcanic margins, were formed at sea level. The shallow ridges are inferred to have promoted widespread deposition and alteration of glassy hyaloclastite—a source of alkalinity. We test this hypothesis by quantifying the prevalence of shallow ridges along Pangea's passive continental margins, and by assessing Neoproterozoic reconstructions of tectonic plates. We find that the most frequently occurring depth range for incipient mid-ocean ridges is 2.1 ± 0.4 km. Ridges with initial elevations of approximately sea level are rare and have anomalous crustal thicknesses >14 km that only occur proximal to large igneous provinces (LIPs). Hyaloclastite is uncommon on mid-ocean ridges as it is generally restricted to water depths of <200 m for tholeiitic basalts, instead forming mostly on intraplate seamounts. Additionally, ocean drilling recently found hyaloclastite to be insignificant along the outer Vøring Plateau (offshore Norway)—an exemplar of a volcanic margin. Reconstructions of Rodinia and associated LIPs demonstrate that volcanic margins potentially hosting minor hyaloclastites were scarce during the late Neoproterozoic. We conclude that the shallow-ridge hypothesis fails to explain the formation of cap carbonates and suggest that other mechanisms such as enhanced continental weathering may be largely responsible.The Neoproterozoic era was punctuated by three significant glaciations. The Sturtian (ca. 717–661 Ma) and the Marinoan (ca. 646–635 Ma) global glaciations were the most severe in Earth's history, with sea ice extending all the way to the equator in a “Snowball Earth” scenario (Hoffman et al., 2017). The Gaskiers glaciation (ca. 580 Ma), in contrast, was a short-lived (≤340 k.y.), mid-latitude, regional event, comparable to Cenozoic glaciations (Pu et al., 2016). Common to all Neoproterozoic glaciations is the global occurrence of cap carbonates—laterally continuous layers of neritic limestone or dolostone immediately overlying glaciogenic deposits or associated erosional surfaces (Grotzinger and James, 2000; Shields, 2005; Hoffman, 2011). The cap carbonates preserve a unique δ13C negative excursion (e.g., Kennedy, 1996; Hoffman et al., 2017) and display unusual textural and compositional features that distinguish them from other carbonate rocks (Kennedy, 1996; Grotzinger and James, 2000; Shields, 2005). Cap carbonate deposition is associated with shallow water supersaturated in carbonate, a warming climate, and continental flooding caused by a eustatic sea-level rise following deglaciation (e.g., Kennedy, 1996; Hoffman et al., 2017). Formation of these puzzling deposits remains controversial, with multiple competing mechanisms proposed concerning the source of the alkalinity and its delivery to the sites of carbonate deposition (Shields, 2005). A novel mechanism put forward by Gernon et al. (2016) and considered plausible by many (e.g., Hoffman et al., 2017; Youbi et al., 2020; Hood et al., 2022) posits that the weathering of hyaloclastite (glassy debris formed by subaqueous magmatic eruptions) along shallow spreading ridges during the break-up of the supercontinent Rodinia supplied alkalinity for cap carbonate deposition via syn-glacial alteration of hyaloclastite to palagonite—a product of volcanic glass hydration (Gernon et al., 2016; Fig. 1). We test the validity of this hypothesis using a quantitative approach to assess the prevalence of anomalously shallow spreading ridges during the break-up of Pangea as a proxy for the break-up of Rodinia. We subsequently use Cryogenian reconstructions of continents, plate boundaries, and large igneous provinces (LIPs) following the break-up of Rodinia to show that the shallow-ridge hypothesis is fundamentally flawed.For our analysis, we first create an updated set of preserved boundaries between stretched continental and ocean crust (COBs) (see the Supplemental Material1 for information on the methodology and global gravity data set used in this study). We choose COB data at 200 km intervals along the COBs and select points 50 km seaward of each COB to sample the age of the ocean crust (Seton et al., 2020; Fig. 2A) and its crustal thickness (Reguzzoni and Sampietro, 2015; Fig. 2B). The distance of 50 km from the interpreted boundary of stretched continental crust was chosen to ensure that we sample crustal thickness based on a grid cell overlying pure ocean crust, given a grid resolution of 0.5 degrees (Reguzzoni and Sampietro, 2015). We then use PyBacktrack (https://github.com/EarthByte/pyBacktrack; Müller et al., 2018) to reconstruct the initial elevation of the seafloor (Fig. 2C) shortly after the onset of seafloor spreading at these points, removing subsequently deposited sediments taken from a global sediment thickness grid (Straume et al., 2019). These initial depths are computed with and without long-term sea-level variations and mantle convection–driven dynamic topography (Young et al., 2022). We consider the difference between modeled dynamic topography at the present day and at the initiation of seafloor spreading (Fig. 2D), using the plate rotation model employed by Young et al. (2022), which is based on the model of Merdith et al. (2021) (see the Supplemental Material).Our assessment of the shallow-ridge hypothesis begins with preserved ocean crust along passive continental margins, where break-up tectonic history is well constrained (Müller et al., 2019). The analysis (Fig. 3) allows us to consider the subsidence history of passive margins that underpin the validity of the shallow-ridge hypothesis (Gernon et al., 2016). Our aim is to test a critical assumption of the hypothesis that the ridge crest in the initial stage of seafloor spreading after break-up is located at sea level (i.e., basement depth is 0) and likely persists at depths of <2 km for 30–35 m.y. after the onset of mid-ocean ridge formation (supplementary information fig. 2 in Gernon et al., 2016).We find that the present-day age of the initial ocean crust ranges from the Late Paleozoic in the eastern Mediterranean to the Neogene along young rifts, with the majority of break-up ages clustered between the early Jurassic (ca. 200 Ma) and the Late Eocene (ca. 35 Ma), reflecting the protracted break-up of Pangea (Fig. 2A). The initial oceanic basement paleo-elevation ranges from a depth of 4 km to shallow subaerial elevation (<0.5 km). The most frequently occurring depths are centered at 2.1 ± 0.4 km (Fig. 3A), which represents an offset toward shallower values from the initial basement depth of 2.6 ± 0.3 km computed by Richards et al. (2018). We primarily attribute this offset to shallowing of the initial paleo-elevation due to subsequent dynamic subsidence along most of Pangea's passive margins (Figs. 2D and 3C; see the Supplemental Material). The tail of shallower-than-expected values (Fig. 3A) includes outliers of initial elevations around sea level and at modest subaerial elevations that are proximal to large-scale mantle upwellings and LIPs (Fig. 2B).The thickness of the initial ocean crust ranges from <1 km where the mantle is exhumed (e.g., the Somali margin; Mortimer et al., 2020) to ~28 km for a small number of sites in regions of plume-related mantle upwelling such as the Iceland plume in the North Atlantic and the Réunion plume, which caused excess volcanism on the western Indian margin (Figs. 2B). The highest proportion of sites occurs within the average range of global thicknesses of ocean crust (Fig. 3C) and agrees with the mean thickness of normal ocean crust (~7 km) distal from anomalous regions such as hot spots (White et al., 1992). The number of sites with crustal thickness >14 km represents the tail end of the distribution, with thicknesses >20 km representing only a very small proportion of all sites (Fig. 3C). These sites typically occur on volcanic passive margins in the vicinity of LIPs (Fig. 2), which are mostly related to mantle-plume activity producing excess volcanism (Coffin and Eldholm, 1994). Despite the abundance of LIPs during the protracted break-up of Pangaea (Fig. 2), very few sites were initially close to sea level, contrary to the prerequisite of the shallow-ridge hypothesis (Fig. 3A; Fig. S1), and these are all associated with extensive LIPs (Fig. 2C).Hyaloclastite formation is considered to be more important above the critical depth of seawater for explosive volcanism, and typically occurs in the final growth stages of seamounts (Staudigel and Schmincke, 1984), rather than on mid-ocean ridges (Bonatti and Harrison, 1988). Tholeiitic magmatic explosivity at mid-ocean ridges is limited to <200 m, and to 1 km for seamount-forming alkalic magmas (Kokelaar, 1986). Rare hyaloclastite deposits have been reported from mid-ocean ridge settings, but these settings are unusual and include fissures on the Eastern Rift Zone of Iceland (Bergh and Sigvaldason, 1991) where the crust has been significantly thickened and uplifted by the Iceland plume, resulting in an anomalously shallow basement depth (Fig. S1), and on the ultraslow-spreading Gakkel Ridge in the Arctic Ocean (Sohn et al., 2008).Although the seafloor is peppered with millions of seamounts, the vast majority are small (<100 m tall), and located on young lithosphere (Wessel et al., 2010). These seamounts are quickly buried by sediment (Wessel, 2007), and in just a few million years subside below the average mid-ocean ridge depth of ~2.6 km as the ocean lithosphere cools, unlikely to ever produce hyaloclastite. Most of the larger seamounts (>1 km tall) were formed in intraplate settings on old ocean crust by hotspot activity (Wessel, 2007), and are common sites for hyaloclastite formation (Batiza, 1982; Staudigel and Schmincke, 1984; Bonatti and Harrison, 1988). However, intraplate seamounts are estimated to comprise a tiny fraction (~0.5%) of all seamounts (Wessel et al., 2010), and only a few are active at any one time (Wessel, 2007).Hyaloclastites can also form in volcanic rifts in early rift environments if rifting occurs below sea level, as documented along the Vøring (offshore Norway) and conjugate Greenland margins (Planke et al., 2000). The transition from rifting to seafloor spreading is commonly marked by an outer volcanic high interpreted from seismic reflection data as a sequence of hyaloclastite flows (Planke et al., 2000). However, recent ocean drilling of the outer high on the Vøring Plateau mainly recovered massive and pillow basalts with only minor hyaloclastite (Planke et al., 2023), implying that the original interpretation based solely on seismic images had significantly overestimated the abundance of hyaloclastite. There is no published evidence that any hyaloclastite deposition continues once seafloor spreading has commenced along volcanic margins, in contrast to the assumption made by Gernon et al. (2016) that hyaloclastite formation continues for up to 35 m.y. after seafloor spreading commences.Although the break-up history of Rodinia is controversial, the plate motion model we used implies a protracted continental break-up lasting ~150 m.y. (Merdith et al., 2021; Fig. 4), comparable to that of Pangea (Müller et al., 2019). Neoproterozoic ocean crust was also similar to modern crust, attaining the average Phanerozoic thickness of 7 km since ca. 0.9 Ga (Moores, 2002). However, unlike the numerous large LIPs associated with the break-up of Pangea, such as the Karoo (southern Africa), High Arctic, and Paraná-Etendeka (South America) LIPs (Fig. 1A) that originally covered between 3.12 Mkm2 and 3.6 Mkm2 (Park et al., 2021), the late Neoproteroic LIPs were far less numerous and extensive, and emplaced in continental interiors. We estimate that Pangea's passive margin length with depths in the range of hyaloclastite formation (0–200 m; Kokelaar, 1986) is ~2000 km—a mere ~1.3% of the total passive margin length of ~156,000 km, and marginally higher (3%) when dynamic topography and sea level are considered. The Franklin LIP (Canadian Arctic; Fig. 4A), emplaced 718 Ma (Dufour et al., 2023), was quite large (2.64 Mkm2), but the Central Iapetus Magmatic Province (CIMP; Fig. 4C) was emplaced in multiple pulses between ca. 615 Ma and 560 Ma over an area ranging from only 0.07–0.36 Mkm2 (Ernst et al., 2021). The absence of continental margins intersecting LIPs during Rodinia's break-up suggests that plume-related volcanic margins and associated dynamic uplift were rare. Consequently, spreading ridges did not form at or near sea level and the potential for any significant hyaloclastite formation via the shallow-ridge mechanism was greatly reduced. This is inadvertently substantiated by most sequences cited by Gernon et al. (2016, see their fig. 2 and their supplementary table 1) in support of the shallow-ridge hypothesis, in which small volumes of hyaloclastites were generated in convergent tectonic settings rather than on passive margins. Additionally, hypersaline and near-freezing bottom water consistent with near-global sea-ice cover (Hoffman et al., 2017) would have considerably slowed down weathering (Coogan and Dosso, 2015) and the alteration of hyaloclastite to palagonite by which alkalinity is supplied to the ocean in the shallow-ridge hypothesis (Fig. 1). Most spreading ridges form at depths of ~2.6 km in the absence of dynamic topography (Fig. 3A) and subside to ~4 km within 35 m.y. (Richards et al., 2018), which casts doubt on extensive hyaloclastite formation for 35 m.y. after the onset of mid-ocean ridge initiation proposed by Gernon et al. (2016).We conclude that the shallow-ridge hypothesis fails to explain the formation of extensive cap carbonates marking the termination of Neoproterozoic glaciations globally. While the genesis of these unique carbonates remains controversial, it is more likely that other mechanisms such as enhanced continental weathering were largely responsible for supplying alkalinity to the Neoproterozoic ocean (Hoffman et al., 2017). It is unlikely that mid-ocean ridge evolution played any significant role in driving cap carbonate formation, even though plate tectonics may have played a role in modulating Cryogenian climate via changes in solid Earth degassing after Rodinia's break-up (Dutkiewicz et al., 2024). Future coupled plate-mantle models for the past billion years will be helpful for elucidating the effect of plate boundary evolution, mantle plumes, and dynamic topography on Rodinia's and Gondwana's passive margins, evolving ocean basins, and ocean chemistry.We thank three anonymous reviewers for their detailed and constructive comments. This research was supported by the Australian Research Council (ARC) Future Fellowship FT190100829 to A. Dutkiewicz and by the National Collaborative Research Infrastructure Strategy (NCRIS) via AuScope.

中文翻译：

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沿浅海脊的海底火山活动并未驱动成冰期盖碳酸盐岩的形成

新元古代“雪球地球”冰川作用的终止在全球范围内的标志是直接覆盖冰川混积岩的横向广泛的浅海帽碳酸盐。这些独特的冰消沉积物的形成需要异常高的碳酸钙饱和度。解释所需海洋碱度来源的一种流行机制是浅脊假说，其中围绕罗迪尼亚碎片的初始扩张脊（假定以火山边缘为主）是在海平面形成的。据推测，浅脊促进了玻璃状透明碎石（碱度来源）的广泛沉积和蚀变。我们通过量化盘古大陆被动大陆边缘浅脊的普遍程度以及评估构造板块的新元古代重建来检验这一假设。我们发现早期洋中脊最常出现的深度范围是 2.1 ± 0.4 公里。初始海拔接近海平面的山脊非常罕见，且异常地壳厚度 >14 公里，仅出现在大型火成岩省 (LIP) 附近。玻璃碎屑岩在洋中脊上并不常见，因为对于拉斑玄武岩来说，它通常仅限于 <200 m 的水深，而主要形成于板内海山。此外，最近的海洋钻探发现，在沃尔令高原（挪威近海）外层（火山边缘的一个典范），透明碎石的含量微不足道。罗迪尼亚和相关唇部的重建表明，在新元古代晚期，可能存在少量玻璃碎屑岩的火山边缘很少。我们得出的结论是，浅脊假说无法解释盖碳酸盐的形成，并表明其他机制（例如增强的大陆风化）可能是主要原因。新元古代时期经历了三个重要的冰川作用。斯图尔蒂安期（约 717–661 Ma）和马里诺安期（约 646–635 Ma）全球冰川作用是地球历史上最严重的，海冰在“雪球地球”情景中一直延伸到赤道（霍夫曼等人，2017）。相比之下，Gaskiers 冰川作用（约 580 Ma）是一次短暂的（≤340 ky）中纬度区域性事件，与新生代冰川作用相当（Pu 等，2016）。所有新元古代冰川作用的共同点是全球范围内出现帽状碳酸盐——直接覆盖冰川沉积物或相关侵蚀面的横向连续的浅海石灰岩或白云岩层（Grotzinger 和 James，2000；Shields，2005；Hoffman，2011）。盖帽碳酸盐岩保留了独特的 δ13C 负偏移（例如，Kennedy，1996；Hoffman 等人，2017），并显示出与其他碳酸盐岩不同的不寻常的结构和成分特征（Kennedy，1996；Grotzinger 和 James，2000；Shields， 2005）。盖帽碳酸盐沉积与碳酸盐过饱和的浅水区、气候变暖、以及冰川消退后海平面上升引起的大陆洪水（例如，Kennedy，1996；Hoffman 等，2017）。这些令人费解的沉积物的形成仍然存在争议，关于碱度的来源及其向碳酸盐沉积地点的输送提出了多种相互竞争的机制（Shields，2005）。 Gernon 等人提出了一种新颖的机制。 （2016）并且被许多人认为合理（例如，Hoffman 等人，2017；Youbi 等人，2020；Hood 等人，2022）认为，玻璃碎石（水下岩浆喷发形成的玻璃碎片）沿着浅层扩散的风化作用在罗迪尼亚超大陆分裂期间，山脊通过将透明碎屑岩转变为火山玻璃水化的产物长菱角石，为盖碳酸盐沉积提供了碱度（Gernon等人，2016年；图1）。我们使用定量方法来测试这一假设的有效性，以评估盘古大陆分裂期间异常浅的扩张脊的普遍程度，作为罗迪尼亚分裂的替代指标。随后，我们在罗迪尼亚解体后使用低温纪重建大陆、板块边界和大型火成岩省 (LIP)，以表明浅脊假说从根本上是有缺陷的。对于我们的分析，我们首先创建一组更新的保留边界拉伸的大陆和海洋地壳 (COB) 之间（有关本研究中使用的方法和全球重力数据集的信息，请参阅补充材料1）。我们沿着COB以200公里的间隔选择COB数据，并选择每个COB向海50公里的点来采样洋壳的年龄（Seton等，2020；图2A）及其地壳厚度（Reguzzoni和Sampietro，2015）图2B）。选择距拉伸大陆地壳解释边界 50 公里的距离，以确保我们基于覆盖纯海洋地壳的网格单元对地壳厚度进行采样，网格分辨率为 0.5 度（Reguzzoni 和 Sampietro，2015）。然后，我们使用 PyBacktrack（https://github.com/EarthByte/pyBacktrack；Müller 等人，2018）在这些点的海底扩张开始后不久重建海底的初始高程（图 2C），随后删除从全球沉积物厚度网格中获取的沉积物（Straume 等人，2019）。这些初始深度是在考虑或不考虑长期海平面变化和地幔对流驱动的动态地形的情况下计算的（Young et al., 2022）。我们使用 Young 等人采用的板块旋转模型，考虑了当今模拟的动态地形与海底扩张开始时的动态地形之间的差异（图 2D）。 （2022），这是基于 Merdith 等人的模型。 （2021）（参见补充材料）。我们对浅脊假说的评估始于沿被动大陆边缘保存的洋壳，那里的分裂构造历史受到很好的限制（Müller 等人，2019）。分析（图二）3）使我们能够考虑被动边缘的沉降历史，这支撑了浅山脊假说的有效性（Gernon等人，2016）。我们的目的是检验以下假设的一个关键假设：海底破裂后扩张初始阶段的脊顶位于海平面（即基底深度为 0），并且可能在 <2 km 的深度持续 30 年。洋中脊形成开始后 –35 米（Gernon 等人，2016 年的补充信息图 2）。我们发现，目前初始洋壳的年龄范围从地中海东部的晚古生代到新近纪沿着年轻的裂谷，大部分分裂年龄集中在早侏罗世（约200Ma）和晚始新世（约35Ma）之间，反映了盘古大陆的长期分裂（图2A）。最初的海洋基底古海拔范围从4公里深度到浅层地面海拔（<0.5公里）。最常出现的深度集中在 2.1 ± 0.4 km（图 3A），这表示从 Richards 等人计算的初始基底深度 2.6 ± 0.3 km 向较浅值的偏移。 （2018）。我们主要将这种偏移归因于随后沿盘古大陆大部分被动边缘的动态沉降导致的初始古海拔变浅（图2D和3C；参见补充材料）。比预期值浅的尾部（图 3A）包括海平面周围初始海拔的异常值以及接近大规模地幔上升流和 LIP 的中等地面海拔（图 2B）。 初始海洋的厚度地壳范围从地幔被挖出的地方 <1 公里（例如，索马里边缘；Mortimer 等人，2020）到与地幔柱相关的地幔上升区域（例如冰岛地幔柱）中的少数地点的约 28 公里。北大西洋和留尼旺羽流，导致印度西部边缘过度火山活动（图 2B）。最高比例的地点出现在全球海洋地壳厚度的平均范围内（图3C），并且与远离热点等异常区域的正常海洋地壳的平均厚度（~7公里）一致（White等人， 1992）。地壳厚度>14 km的站点数量代表了分布的尾端，厚度>20 km的站点只占所有站点的很小一部分（图3C）。这些地点通常出现在 LIP 附近的火山被动边缘上（图 2），这主要与产生过量火山活动的地幔柱活动有关（Coffin 和 Eldholm，1994）。尽管在盘古大陆长期分裂过程中存在丰富的 LIP（图 2），但最初接近海平面的地点很少，这与浅脊假说的先决条件相反（图 3A；图 S1），并且这些都与广泛的 LIP 有关（图 2C）。在海水爆发性火山活动的临界深度以上，玻璃碎石的形成被认为更为重要，通常发生在海山的最后生长阶段（Staudigel 和 Schmincke，1984），而不是发生在洋中脊（Bonatti 和 Harrison，1988）。大洋中脊的拉斑斑岩岩浆爆发力限制在<200 m，对于形成海山的碱性岩浆则限制在1 km（Kokelaar，1986）。据报道，在大洋中脊环境中发现了稀有的透明碎石矿床，但这些环境并不常见，包括冰岛东部裂谷带的裂缝（Bergh 和 Sigvaldason，1991），那里的地壳因冰岛羽流而显着增厚和抬升，导致位于异常浅的基底深度（图 S1），以及北冰洋超慢速扩张的 Gakkel 海脊（Sohn 等，2008）。虽然海底布满了数百万个海山，但绝大多数都是小型的（< 100 m 高），位于年轻岩石圈上（Wessel 等，2010）。这些海山很快被沉积物掩埋（Wessel，2007），并且在短短几百万年之内，随着海洋岩石圈的冷却，沉降到平均洋中脊深度约 2.6 公里以下，不太可能产生玻璃碎屑岩。大多数较大的海山（>1 公里高）是由热点活动在旧洋壳的板内环境中形成的（Wessel，2007 年），并且是透明碎石形成的常见地点（Batiza，1982 年；Staudigel 和 Schmincke，1984 年；Bonatti 和 Harrison） ，1988）。然而，据估计，板内海山仅占所有海山的一小部分（~0.5%）（Wessel et al., 2010），并且在任何时间只有少数海山处于活动状态（Wessel, 2007）。玻璃碎屑岩也可以在火山岩中形成如果裂谷发生在海平面以下，则早期裂谷环境中会出现裂谷，正如沿 Vøring（挪威近海）和共轭格陵兰边缘记录的那样（Planke 等人，2000 年）。从裂谷到海底扩张的过渡通常以外部火山高地为标志，根据地震反射数据将其解释为一系列玻璃碎屑流（Planke et al., 2000）。然而，最近对沃尔令高原外高处的海洋钻探主要发现了块状玄武岩和枕状玄武岩，仅含有少量的透明碎石（Planke等，2023），这意味着最初仅基于地震图像的解释明显高估了透明碎石的丰度。与 Gernon 等人的假设相反，没有已发表的证据表明一旦海底沿着火山边缘开始扩张，任何透明碎石的沉积就会继续。 （2016）认为，海底扩张开始后，玻璃碎屑岩的形成持续长达 35 my。尽管罗迪尼亚的分裂历史存在争议，但我们使用的板块运动模型表明，大陆分裂持续约 150 my（Merdith 等人，2016）。 ，2021；图 4），与盘古大陆相当（Müller 等人，2019）。新元古代洋壳也与现代地壳相似，显生宙自大约 1997 年以来平均厚度达到 7 公里。 0.9 Ga（摩尔斯，2002）。然而，与盘古大陆分裂相关的众多大型 LIP 不同，例如卡鲁（南部非洲）、高纬度北极和巴拉那-埃滕德卡（南美洲）LIP（图 1A），它们最初覆盖的面积在 3.12 Mkm2 和 3.6 Mkm2 之间（ Park 等人，2021）认为，新元古代晚期的 LIP 数量和范围要少得多，而且分布在大陆内部。我们估计，盘古大陆的被动边缘长度（深度在玻璃碎屑岩形成范围内（0-200 m；Kokelaar，1986）约为 2000 公里，仅占被动边缘总长度约 156,000 公里的约 1.3%，并且略高一些（ 3%）当考虑动态地形和海平面时。富兰克林 LIP（加拿大北极；图 4A）在 718 Ma 就位（Dufour 等，2023），相当大（2.64 Mkm2），但中土卫八岩浆省（CIMP；图 4C）在多个脉冲中就位大约之间615 Ma 和 560 Ma 的区域范围仅为 0.07–0.36 Mkm2（Ernst 等人，2021）。在罗迪尼亚分裂期间，没有大陆边缘与 LIP 相交，这表明与火山喷发相关的火山边缘和相关的动态隆起很少见。因此，扩张脊不会在海平面或海平面附近形成，并且通过浅脊机制形成任何重要的透明碎屑岩的可能性大大降低。 Gernon 等人引用的大多数序列无意中证实了这一点。 （2016，参见他们的图 2 和补充表 1）支持浅山脊假说，其中少量的透明碎屑岩是在聚合构造环境中生成的，而不是在被动边缘生成的。此外，与近乎全球海冰覆盖相一致的高盐度和接近冰冻的底层水（Hoffman 等人，2017 年）将大大减缓风化（Coogan 和 Dosso，2015 年）以及透明碎屑岩向长角长石的转变，从而降低碱度。在浅海脊假说中向海洋供应水（图 1）。在没有动态地形的情况下，大多数扩张脊形成于约 2.6 公里的深度（图 3A），并在 35 米内沉降到约 4 公里（Richards 等人，2018），这对 35 米后广泛的玻璃碎屑岩形成产生了怀疑。 Gernon 等人提出的洋中脊起始的开始。 (2016)。我们的结论是，浅山脊假说无法解释标志着全球新元古代冰川作用终止的大范围碳酸盐岩帽的形成。虽然这些独特碳酸盐的成因仍然存在争议，但更有可能的是，其他机制（例如增强的大陆风化）在很大程度上为新元古代海洋提供了碱度（Hoffman et al., 2017）。尽管板块构造可能通过罗迪尼亚解体后固体地球脱气的变化在调节成冰纪气候方面发挥了作用，但大洋中脊的演化不太可能在驱动盖碳酸盐岩形成方面发挥任何重要作用（Dutkiewicz 等，2024） ）。过去十亿年的未来板块-地幔耦合模型将有助于阐明板块边界演化、地幔柱和动态地形对罗迪尼亚和冈瓦纳被动边缘、演化海洋盆地和海洋化学的影响。我们感谢三位匿名审稿人他们详细而有建设性的意见。这项研究得到了澳大利亚研究委员会 (ARC) 未来奖学金 FT190100829 授予 A. Dutkiewicz 的支持，并通过 AuScope 得到了国家合作研究基础设施战略 (NCRIS) 的支持。