The Relationship between the Time Difference of Formation Water Infiltration Rate, Tectonic Movement, and the Formation Pressure Coefficient

Abstract

The study of formation pressure holds great significance for both exploration and development. The formation pressure coefficient is a crucial parameter in geology, encompassing various aspects. Numerous models exist to explore its influencing factors, yet they remain highly controversial. Our research has gathered data on formation pressure states and tectonic movement rates from dozens of large sedimentary basins worldwide. We delved into the patterns of formation pressure changes during basin deposition, subsidence, and tectonic uplift, taking into account the permeability of formation water. The findings reveal that during the Neogene and beyond, rapid deposition or tectonic uplift can cause the infiltration of formation water to lag behind tectonic movements, resulting in overpressure. Conversely, if recent tectonic movements are slow, formation water will complete its infiltration process ahead of tectonic changes, bringing the formation pressure to a hydrostatic state. Consequently, we have concluded that abnormal formation pressure primarily depends on the rate of tectonic movements during the Neogene and Quaternary periods. This study also proposes four formation pressure models, paving the way for a comprehensive understanding of formation pressure within a unified theoretical framework.

1. Introduction

Stratigraphic pressure is an extremely important physical quantity in geological research, with overpressure and ultrahigh pressure receiving particular attention. This phenomenon is observed in approximately two-thirds of sedimentary basins worldwide [1] and in over half of China’s hydrocarbon-producing layers [2]. Given its close association with hydrocarbon distribution, overpressure has become a pivotal aspect of research in hydrocarbon accumulation dynamics [3,4,5,6,7]. In China, formation pressure is commonly classified as ultra-low pressure (pressure coefficient <0.75), low pressure (0.75–0.9), normal pressure (0.9–1.2), high pressure (1.2–1.5), and ultra-high pressure (>1.5) [8]. The effect of overpressure on oil and gas is multifaceted. It not only influences the hydrocarbon generation process of organic matter and the occurrence state of hydrocarbons but also concerns the dynamics and direction of migration. In addition, overpressure plays a significant role in the preservation and destruction of gas reservoirs and has a profound impact on the spatial distribution of oil and gas reservoirs, such as episodic hydrocarbon expulsion [9], episodic charging and episodic reservoir formation [10], and overpressure migration [11]. Moreover, overpressure can contribute to high hydrocarbon production [12]. Meissner (1987) [13] proposed a pressure cycle controlled by hydrocarbon generation, transitioning from initial normal hydrostatic pressure to abnormal high pressure (due to hydrocarbon generation and expansion), then to abnormal low pressure (as migration halts and hydrocarbon dissipation exceeds supply), and finally back to normal pressure. Additionally, overpressure implications extend to the safety of drilling and geological engineering operations and the development planning of oil and gas fields [14,15]. Therefore, the distribution and genesis mechanism of overpressure have always been of great concern to geologists and exploration workers, and the significance of research in this area is profound.

The causes of overpressure have been widely discussed, with undercompaction, hydrocarbon generation, and tectonic compression being recognized as the three common pressure-increasing mechanisms [16,17,18,19,20]. Wang et al. (2019) [21] proposed that tectonic compression can reduce porosity, thereby generating overpressure. The two most common causes of abnormally high pressure in petroleum basins are compaction disequilibrium and hydrocarbon generation [22]. Osborn and Swarbrick (1997) [23] suggested that unbalanced compaction is a viable mechanism for overpressure during rapid burial of thick mudstone sequences while dismissing other mechanisms, sparking significant debate. Bethke (1986) [24] quantitatively analyzed the efficiency of dehydration during the transformation of smectite to illite on formation pressure, concluding that its effect is negligible. Both undercompaction and hydrocarbon generation are considered the primary factors leading to overpressure. Qiu et al. (2020) [25] argued that most factors causing abnormal pressure occur only under specific conditions, with undercompaction, hydrocarbon generation, and tectonic stress remaining the most common controlling factors. Due to the influence of overpressure genesis, there can be a negative effect on the generation of hydrocarbons. For instance, the overall impact of overpressure on the thermal evolution of organic matter can be divided into four levels, but generally, it has an inhibitory effect, albeit with varying degrees of inhibition [26].

In summary, there are various factors that can cause formation overpressure. Currently, there is no single mechanism that can explain overpressure under all conditions. For example, undercompaction can only explain overpressure during the sedimentation process but not during uplift. Tectonic compression can explain overpressure in foreland basins but cannot explain overpressure in extensional fault basins. Hydrocarbon generation and pressurization can cause overpressure, but they cannot explain overpressure in the absence of oil and gas. For example, the Kela 2 gas field has overpressure, but Kela 1 and Kela 3, which do not contain oil and gas, still exhibit ultra-high pressure. Each hypothesis can only explain a specific type of overpressure, indicating that these piecemeal, customized, or patchwork theories of overpressure genesis may inherently have some flaws. Generally speaking, the more geological factors listed, the further away we may be from the truth. Whether there is a unified, truly dominant factor that can explain formation pressure under different conditions is the goal of this article to explore in depth.

A common, yet often overlooked factor, is formation water, which plays a pivotal role in underground geological fluids. It is a crucial element throughout the entire process from the initial development of sedimentary basins to the advanced diagenesis stage and should be considered the dominant force. However, most scholars believe that the osmotic effect of formation water on formation pressure is less efficient in pressure increase compared to other factors, leading to a general disregard for its impact [27]. Neuzil (1995) [28] demonstrated unbalanced abnormal pressure phenomena from the perspective of formation water dynamics, controlled by compaction (with a high deposition rate of >200 m/Ma), diagenesis, and deformation, independent of hydrocarbon generation. Nevertheless, this has not garnered sufficient attention. Indeed, due to the small radius of water molecules, groundwater in most lithologic formations transitions relatively quickly from an abnormal pressure state to a more balanced hydrostatic pressure state compared to the timescale of tectonic movements. However, in well-sealed mudstones, shales, and gypsum–salt layers, there is a certain lag in water infiltration when formation pressure changes, and equilibrium is not achieved instantaneously or within a short period relative to tectonic movements. Therefore, exploring the nature of formation overpressure from the perspective of formation water is of great significance, as it may help us uncover the truth.

This article focuses on the changes of formation water before and after diagenesis and attempts to explore the relationship between abnormal pressure variations and formation water. By combining the permeability of formation water, four models of formation pressure evolution are proposed, and an analysis of major basins both domestically and internationally reveals their universality. It is also suggested that formation water has been fully discharged through the seepage process before diagenesis, eliminating the possibility of undercompaction due to inadequate drainage. After diagenesis, formation water completes the transition from overpressure to normal pressure through permeation, and this switch does not result in a significant increase in porosity.

2. Data and Methodology

The data utilized in this study comprises basin formation pressure states, tectonic evolution histories, and tectonic movement rates from over 50 basins or regions both domestically and internationally. Additionally, measured data on organic and inorganic pores from 163 overpressured strata in 11 basins or regions were included, along with calculations of their respective proportions.

The method of study is to analyze the pressure changes of stratigraphic water in the formation of sedimentary basins and the process of diagenetic evolution. In porous media, the water is smaller than other hydrocarbon molecules, and 3.2 Å (1 Å = 10–10 m), so there is water penetration in places with abnormal stress. The following

Table 1. Effective of different molecules (Unit: Å).

Molecule	Effective Diameter Å	Molecule	Effective Diameter Å
H2O	3.2	C3H8	5.1
He	2.0	Benzene	4.7
CO2	3.3	Normal alkanes	width: 4.2 length: 4.2~4.8
N2	3.4	Cyclohexane	5.4
CH4	3.8	Complex cyclic compounds	15~20
C2H6	4.4	Bitumen	>50

shows the effective diameter of different molecules [29]. Under the conditions of formation, the deeper the pressure, the local if there is a condition that exceeds the hydrostatic force condition, the water has the tendency to permeate and tend to the hydrostatic pressure.

Although water has strong permeability, there is a certain lag time in strata of different lithologies, and the process is not instantaneously completed. According to the Poiseuille formula [30,31], assuming a mudstone stratum thickness of 200 m and a pore radius of 2 nm, it only takes 1.5 Ma for the formation water in the reservoir to drop from an initial 2.6 times overpressure to 1.2 times overpressure due to the permeability of the mudstone, as shown in

Table 2. Timetable for overpressure relief of mudstone with different pore sizes and thicknesses.

Pore Size	Thickness (m)
(nm)	50	100	200	500	1000	2000
Pressure relief to 1.6 times:
1	0.297	1.189	4.756	29.723	-	-
2	0.074	0.297	1.06	7.43	29.723	-
5	0.012	0.048	0.190	1.189	4.756	19.023
10	0.003	0.012	0.048	0.297	1.189	4.756
20	0.001	0.003	0.012	0.074	0.297	1.189
50	0	0	0.002	0.012	0.048	0.190
Pressure relief to 1.2 times:
1	0.376	1.506	6.022	37.640	-	-
2	0.094	0.376	1.506	9.410	37.640	-
5	0.015	0.060	0.241	1.506	6.022	24.090
10	0.004	0.015	0.06	0.376	1.506	6.022
20	0.001	0.004	0.015	0.094	0.376	1.506
50	0	0	0.002	0.015	0.06	0.241

. Wu (2024) [32] studied the pore size distribution of 7 shale samples from the Shahejie Formation in the Jiyang Depression and found that pores ranging from 0–4 nm accounted for 59%, indicating that this pore radius is reasonable. During rapid accumulation, the initial condition is that both the stratum and stratum framework bear the crustal stress together, so 2.6 is taken. This calculation can at least illustrate that the permeability of water in the stratum is relatively fast.

During the syngenetic and early diagenetic stages, due to higher porosity and permeability coefficients, the formation pressure remains at hydrostatic pressure (normal pressure). The subsequent analysis in this article mainly focuses on strata in the middle and late diagenetic stages.

A set of regional cap rocks with certain sealing properties may contain multiple lithologies, such as thick shale, gypsum–salt rock, and thin sand layers. To maintain overpressure in the formation itself or its underlying sandstone reservoir, it is necessary to calculate the time required for pressure relief. The calculation can be approximated by the following formula:

$$t=A{\displaystyle \sum _{i=1}^n\frac{S_S{h}_i^2}{K_i}}$$

(1)

In the formula, t represents the required time (s), h_i is the thickness of each sublayer (m), and K_i is the permeability coefficient of each sublayer (m/s). The Ss represents the water storability (1/m). A represents the dissipation coefficient, which can be taken as 1.5. Moreover, n represents the number of layers.

According to Li et al. (2024), the Ss can be taken as 0.0008 [33]. Based on the research by Yu et al. (2008) [34], the permeability coefficient of mudstone caprock can be taken as 4 × 10⁻¹² m/s. Calculating separately for caprock thicknesses of 100 m, 200 m, 300 m, and 500 m, the required time for abnormal pressure release in the strata is approximately 0.23 Ma, 0.9 Ma, 2.1 Ma, and 5.71 Ma. These calculations are indicative and show that the process of formation water permeation is not instantaneous but relatively fast compared to geological timescales.

3. Two Modes of Forming Overpressure

Based on the characteristics of formation water and its relationship with the rate of tectonic movement, two potential modes for the formation of overpressure have been proposed. The first mode involves rapid subsidence leading to overpressure, as exemplified by the Nanpu Sag. The second mode is associated with rapid tectonic uplift, leading to overpressure, as observed in the Sichuan Basin.

3.1. Active Overpressureing (Overpressure Formed by Rapid Subsidence)

Statistics show that overpressure is prone to occur in strata at the mesodiagenetic and telodiagenetic stages when there is rapid deposition and subsidence in the basin during recent geological periods, such as the Neogene, especially the Quaternary. The following three conditions need to be met: (1) The strata where overpressure occurs require a higher degree of diagenesis, ranging from the mesodiagenetic to the telodiagenetic stage. (2) The rapid subsidence should have occurred relatively recently, during the Neogene or Quaternary; more distant rapid subsidence, such as in the Paleogene or earlier strata, cannot form overpressure. (3) The current depth of deposition and burial exceeds the maximum burial depth in its geological history. At this time, compaction, cementation, and other diagenetic processes are still continuing, pores are decreasing, pore water is pressurized, and there is not enough time for pressure relief. The strata are still in the process of deposition and subsidence, leading to the formation of overpressure. This process is called active pressurization, or rapid subsidence overpressure, i.e., deposition and subsidence occur faster than the permeation rate of formation water.

The widespread existence of overpressure in the Shahejie Formation of the Nanpu Sag, with a maximum overpressure of up to 1.6 times the normal pressure [35], is shown in Figure 1. Weak overpressure is observed in the second and third members of the Dongying Formation, while overpressure occurs in the first member of the Shahejie Formation. The second and third members of the Shahejie Formation exhibit abnormally high pressure. The pressure distribution within the formation does not completely align with the formation trend and can cross different geological times and layers.

The following is an explanation of the reasons for overpressure in the Nanpu Sag using the high permeability of formation water as an example. Recently, two sets of strata have been rapidly deposited. One is the Late Miocene Guantao Formation with a maximum thickness of 1 km, and the other is the Pliocene Minghuazhen Formation with a thickness of 2 km. This has resulted in a time lag between the permeation of formation water and the speed of tectonic movement. In the late rifting stage of the Nanpu Sag, towards the end of the Dongying Formation deposition, faulting activities ceased, resulting in a relatively long period of tectonic stability. During this time, the third member of the Shahejie Formation (Sha-3) was buried at a depth of 3 km, and the formation pressure was hydrostatic, allowing for the completion of the permeation process of formation water. Subsequently, during the depression period, the 1-kilometer-thick Guantao Formation was deposited at a relatively rapid rate, deepening the Sha-3 to a depth of 4 km. The Shahejie Formation had largely completed mechanical compaction and entered the early stage of the middle diagenesis phase A1 [36]. During this stage, pores continued to shrink, and the increase in overburden load led to a situation where both the fluid and the rock matrix shared the total overburden pressure.

To describe this process, we borrow the idea from McKenzie’s (1978) instantaneous tensile model [37]. We decompose the sedimentary process into an instantaneous deposition to more concisely analyze the stabilization of its temperature and pressure fields. As shown in Figure 1, it is assumed that at the end of the Dongying Formation deposition, the third member of Shahejie Formation is at a depth of 3 km, which is in the transition period of fault depression. At this time, the formation pressure is a hydrostatic pressure of 30 MPa. Afterwards, 1 km of the late Miocene Guantao Formation was deposited, and the load of the overlying strata will be applied to the pore fluid and skeleton of the third member of Shahejie Formation. The initial formation pressure is calculated as 30 + 25 MPa. However, the overpressure borne by the pores will decay with time, changing to 30 + 10 MPa. During the deposition of the Minghuazhen Formation, due to the short time period, it can be qualitatively assumed that its formation pressure will remain at 80% (represented by a, where a = 0.8). Therefore, the formation pressure at this time is 30 + [10 + (25 − 10) × 0.8] = 52 MPa.

Subsequently, during the Pliocene, the Minghuazhen Formation with a thickness of 2 km was deposited, adding another 2 × 25 MPa of overlying pressure. This pressure also acts on the pore fluid. Since the Minghuazhen Formation dates back approximately 2.1 million years, which is a relatively short period of time, the pore water pressure has not had time to dissipate. Therefore, the final formation pressure is calculated as 52 + 50 = 102 MPa. At this point, the depth is 6 km, and when converted into a pressure coefficient c, it is 1.67, which is basically consistent with actual measurements. Although the Dongying Formation developed lacustrine mudstone, due to the long exposure time of its top surface and the development of coarse-grained sedimentary facies such as braided rivers and meandering rivers in the Neogene Guantao and Minghuazhen Formations, its formation pressure can quickly permeate and form normal pressure. However, the abnormally high-pressure formation water in the Shahejie Formation needs to pass through the thick mudstone layer of the Dongying Formation to completely relieve pressure (permeate). Therefore, in the short term, the Shahejie Formation can still maintain abnormally high pressure. That is, the thicker the covering layer and the better the sealing layer, the longer the time to maintain abnormal pressure, resulting in a higher pressure coefficient as one goes deeper (Figure 1).

According to the aforementioned condition (2), during earlier periods such as the Paleozoic and Mesozoic, even if there was rapid deposition and subsidence at a certain stage, high formation pressure would be relieved due to the long duration of time.

In summary, the time difference between the permeation of formation water and tectonic movement can provide a good semi-quantitative explanation for the evolution of formation pressure.

3.2. Passive Pressurization (Overpressure Formed by Rapid Tectonic Uplift)

For strata in the mesodiagenetic and telodiagenetic stages, overpressure can also form after recent rapid tectonic uplift, known as passive pressurization (or passive overpressure). This is because strata with higher diagenesis have poor permeability, and pore fluid pressure will be maintained for a period of time after uplift. The cementation process surrounds the fluid and slows down the pressure release process. Although the uplift process may cause some pressure loss due to pore rebound and cooling, it is an order of magnitude smaller compared to the new pressure coefficient c converted at a shallower depth. Passive pressurization also needs to follow the aforementioned conditions (1) and (2).

The mainstream view of most scholars currently is that uplift should result in a low-pressure trend. Their main argument is that after uplift, temperature decreases, pressure drops, and rebounding increases pores, leading to an increase in the available space for fluid volume within the pores. According to the formula of state for liquid water, formation pressure would then decrease [38]. However, these viewpoints do not adequately consider the permeation process of formation water. Both the Wufeng–Longmaxi shale formation and the Triassic Xujiahe formation in the Sichuan Basin commonly exhibit high overpressure, as shown in Figure 2. Essentially, this is the result of recent (Pliocene) rapid tectonic uplift exceeding the rate of water permeation process completion. Similar to the aforementioned active overpressure, the closer the tectonic uplift occurred to the present day, the better, especially during the Neogene and Quaternary periods.

According to Deng et al. (2009) [39], their research on the restoration of surface denudation in the Sichuan Basin reveals that the basin experienced two intense uplifts during the Cenozoic Era. Specifically, the Paleogene uplift averaged approximately 4000 m, while the Neogene uplift was about 3000 m. The target layer, the Xujiahe Formation of the Upper Triassic, currently has a burial depth of approximately 4 km. Assuming an initial state at the end of the Cretaceous, where the formation was under normal pressure with a burial depth of 11 km and an initial formation pressure of 110 MPa. This formation pressure is assumed to be the result of long-term permeation of pore water, forming a normal hydrostatic pressure due to decompression, independent of the effective stress on the framework. After the Paleogene uplift of 4 km, the depth decreased to 7 km. If this process had occurred over a prolonged period, the pressure would have theoretically dropped to a hydrostatic pressure of 70 MPa due to decompression. However, water permeation has a certain lag and would not have completely reduced to hydrostatic pressure. Additionally, the reduction in overlying strata weight only serves to “decompress” the formation matrix. The effective stress (supporting stress) experienced by the formation matrix (framework) should be significantly higher than 110 MPa. Assuming a matrix density of 2.6 g/cm³, the effective stress on the matrix would be approximately 286 MPa (calculated as 11 km × 26 MPa/km = 286 MPa). Moreover, the overpressured stratum has already lithified. Temporarily disregarding factors such as uplift rebound of stratum thickness, pore enlargement, and pressure reduction due to cooling (their magnitudes will be estimated in later chapters), the pore pressure remains essentially unchanged. A reduction in the weight of the overlying stratum can only decrease the skeletal support stress, while the pressure within the pores can still be maintained at 110 MPa. Only through a prolonged period of pressure relief can it slowly decrease to the normal hydrostatic pressure of 70 MPa, corresponding to a depth of 7 km. At the end of the Paleogene, this pressure coefficient had already increased to 1.57 due to uplift.

As evolution progressed to the Neogene period, pressure attenuation needs to be considered. At this point, based on the length of burial time, similar to what was mentioned earlier, a qualitative formation pressure retention coefficient a = 0.8 is assigned. This reduces the pressure coefficient from 110 MPa to 88 MPa. During the Neogene period, there was a further uplift of 3 km, changing the depth from 7 km to 4 km. Similarly, due to the relatively short time since then, it can be considered as having no attenuation. Therefore, after two uplifts and based on the new depth of 4 km, the pressure coefficient c increases to 2.2. Studies by multiple scholars have indicated that the uplift during the Neogene period occurred relatively recently, likely during the Late Cenozoic era. Electron spin resonance dating has determined that the oldest age of the Dayi conglomerate in the western Sichuan Basin is between 2.6 and 3.6 Ma. Many regions along the northern edge of the Tibetan Plateau have also deposited a similar set of coarse-grained sediments, such as the Xiyu conglomerate and Jishi conglomerate, with paleomagnetic and fission track ages dating back to the Late Cenozoic (2.5 to 3.5 Ma). This reflects that the latest period of intense uplift in the Tibetan Plateau and Sichuan Basin ended around 2.5 Ma [40]. Compared to other geological timescales, this period is relatively short, and the pore fluid pressure has not had sufficient time to dissipate. Figure 3 shows the burial history of the Sichuan Basin, indicating that the era of the latest rapid uplift and denudation is relatively recent.

If it is a recent rapid uplift, the following approximate formula can be used to calculate the uplift and denudation thickness Δh:

Δh = p × 100 − z_pmax

(2)

where p represents the formation pressure at the depth z_pmax, which corresponds to the maximum formation pressure coefficient, measured in MPa. Using this value for calculations is more accurate as other well sections, such as shallower depth intervals, may rapidly decompress during uplift and become distorted. This formula is only an approximation, as it may involve multiple recent episodes of uplift.

This model requires that the faster the uplift speed and the closer the time of uplift to the present, the better. The closely spaced JY1, PY1, and LY1 wells (<50 km) in the Wulong area on the southeastern edge of the Sichuan Basin all began to uplift during the Cretaceous period. However, JY1 experienced rapid uplift in a relatively recent time, resulting in a 1.5-fold passive overpressure in the Wufeng–Longmaxi shale formation. In contrast, PY1 and LY1 underwent uniform and slow uplift, leading to normal or low pressure, fully validating this reasoning. Most shale oil and gas basins in the United States are characterized by normal to low pressure, which is typically associated with long-term slow uplift and denudation processes that began a considerable time ago. For example, the New Albany shale in the eastern part of the Illinois Basin has been undergoing tectonic uplift since the early Triassic period with a pressure coefficient of 0.98. The Antrim shale in the Michigan Basin has experienced prolonged uplift following the deposition of Jurassic strata, resulting in a pressure coefficient of 0.4–0.6. The Barnett shale formation in Eastland County, Fort Worth Basin, has been uplifting since the end of the Cretaceous period K1 with a pressure coefficient of 1.0. Relatively speaking, the tectonic uplift of the Eagle Ford shale formation in South Texas occurred more recently, starting with rapid uplift at 15 Ma during the Miocene, resulting in a pressure coefficient of 1.27 and weak overpressure, forming a phenomenon of rapid uplift and passive overpressure.

At present, there is no unified understanding of the causes of overpressure. Relatively recognized causes include hydrocarbon generation and pressure boosting, tectonic compression, hydrothermal pressure boosting, undercompaction, pore rebound, and temperature reduction after uplift. These mentioned phenomena may all exist, but compared to the recent tectonic changes proposed in this paper, the impact on formation pressure is 1–2 orders of magnitude smaller.

4. Formation of Normal or Low Pressure Patterns

4.1. Normal Pressure Pattern

When tectonic movement is slower than the permeation rate of formation water, normal pressure will occur, referred to here as the long-term static normal pressure pattern. Due to slow tectonic uplift, subsidence, or stagnant tectonic activity, formation water has sufficient time to leak, resulting in normal pressure.

Since the beginning of the Cenozoic era, the Songliao Basin has maintained a tectonically stable (dormant) static phase for an extended period, in a state of normal to slightly elevated pressure [41].

The tectonic activity in the Erlian Basin stagnated during the Late Cretaceous and maintained a very slow uplift for 100 Ma (Figure 4). As a result, the Late Cretaceous strata exposed on the surface did not consolidate into rock and remained in a long-term static dormant state, as shown in the figure below (modified from Chen et al. (2015)) [42]. Xu et al. (2007), in explaining the normal pressure in the Erlian Basin, used analogies with the Yinggehai Basin and the Bohai Bay Basin. They employed the classic anomaly of a sudden increase in mudstone porosity at depths of 1500–2500 m to argue that the absence of a high-porosity zone in this depth range in the Erlian Basin indicates the absence of overpressure [43]. However, they overlooked the fact that many areas in the second member of the Tengger Formation in the Erlian Basin have been eroded by nearly 1000–2000 m [44], and there is artificiality in the porosity fitting of the Yinggehai Basin. Thus, the use of undercompaction to explain the cause of overpressure is clearly unsuccessful.

4.2. Large Uplift and Small Subsidence Low Pressure Pattern

Large uplift and small subsidence refer to a long-term significant uplift followed by a rapid but small-scale subsidence that does not reach the maximum burial depth in history. Taking the Sulige gas field as an example, it is an abnormally low-pressure gas field. Since the Late Cretaceous (about 95 Ma), it has been in a state of uplift and denudation for an extended period, with a relatively slow denudation rate, allowing sufficient time for the internal overpressure to release. Another possible reason is that the Mesozoic strata in the northern Ordos region have relatively coarse lithology, with no development of source rocks and poor capping layers [45]. The sealing ability of the overlying strata is poor, and the mudstone layers have thinner individual and cumulative thicknesses, leading to rapid water permeation and pressure relief, resulting in low pressure.

A more important reason may be that the ancient uplifted terrain was higher than it is now. According to the research by Zhao et al. (2009), the Weihe, Yinchuan, and Hetao fault depressions formed successively at the end of the Paleocene to Eocene, and their timing and tectonic patterns are no different from those of the Bohai Bay and other basins in the east. The Shanxi Graben System developed later [46]. Li et al. (2012) concluded that the Sulige gas field experienced a small-scale sedimentary subsidence during the Neogene [47]. This means that since the Cenozoic era, there have been fluctuations rather than continuous uplift, resulting in a possible uplift to a higher elevation than present after the Late Cretaceous, followed by a long period of static conditions. After releasing pressure, hydrostatic pressure (normal pressure) was formed, as shown in Figure 5. The shallowest top surface of the P₂s in the Shihezi Formation is 3120 m; during the Neogene, there was a small-scale subsidence, and the P₂s lowered to 3210 m, not having enough time to adjust by “replenishing water” to the dense He 8 reservoir section and not yet returning to normal hydrostatic pressure, forming the present abnormally low-pressure gas reservoir. In the aforementioned active overpressure model, when a new sedimentary load is added to the original equilibrium state, there is a tendency for pore space to shrink, allowing fluids within the pores to share the load of the overlying strata. However, this does not occur during small-scale subsidence after uplift. Because the latest subsidence did not reach the maximum depth in history (the P₂s lowered to 4150 m at the end of the Early Cretaceous, 97 Ma), the organic matter maturity and porosity would remain at their maximum states, and the diagenesis process would stagnate. The fluids within the pores would not share the additional overlying load with the skeleton until the latest subsidence exceeds the original maximum depth, at which point compaction and increased maturity and other diagenesis processes would resume. Therefore, the numerical deviation of abnormal low pressure from normal pressure can be used to qualitatively calculate the relative height of the highest uplift during geological history compared to the present. The formula is as follows:

Δh = 100(P_j − P) = 100(z_j/100 − P) = z_j(1 − c)

(3)

where Δh represents the uplift height in meters (m), P represents the formation pressure at the measuring point in megapascals (MPa), Pj represents the hydrostatic pressure at the corresponding depth z_j of the measuring point in megapascals (MPa), z_j represents the depth of the measuring point in meters (m), and c represents the pressure coefficient.

The pressure coefficient of the Sulige gas reservoir is relatively low (with an average of 0.85), and the main layer has a deep burial depth (with an average of 4000 m). Based on these calculations, its uplift height is approximately 600 m higher than it is now.

In summary, the formation of overpressure may be influenced by factors such as hydrocarbon generation, aquathermal pressurization, and undercompaction, but these factors are not determinative. Instead, overpressure is primarily caused when the rates of uplift and subsidence exceed the permeation rate of formation water. The impact of overpressure on hydrocarbon reservoirs should also be adjusted accordingly. For instance, overpressured compartments can facilitate the accumulation of hydrocarbons, promote hydrocarbon expulsion, episodic hydrocarbon expulsion, episodic charging, and episodic accumulation. These processes can be summarized into the following four models (Table 2 and Figure 6):

(1)

The active overpressure model of rapid subsidence is known as the Nanpu Model, as shown in Figure 6a.

(2)

The passive overpressure model of rapid tectonic uplift is known as the Sichuan Model, as shown in Figure 6c.

(3)

The long-term static normal pressure model is known as the Songliao Model, as shown in Figure 6b.

(4)

The model of significant uplift followed by minor subsidence and low pressure is known as the Sulige Model (low pressure), as shown in Figure 6d.

These four models can determine the formation pressure state within a unified theoretical framework, eliminating the need to adopt different hypotheses for formation pressure in different situations (tension, compression, uplift, subsidence, and with or without oil and gas involvement) (

Table 3. Four models of formation pressure conditions.

Number	Models	Factor	Characteristics of Tectonic Movement	Pressure State	Examples
1	Nanpu Model	Active overpressure	Rapid subsidence	Overpressure	Nanpu Oilfield
					Qaidam Basin
2	Sichuan Model	Passive overpressure	Rapid uplift and denudation	Overpressure	Sichuan Basin
					Kuqa Depression
					Junggar Basin
3	Songliao Model	Long-term static	Slow uplift	Normal pressure	Songliao Basin
			Slow settling		Erlian Basin
			Neither rise nor fall		Baise Basin
4	Sulige Model	Large uplift and small subsidence	Large uplift + small settlement	Abnormally low pressure	Sulige Gas Field

).

5. Verification of Universality

To demonstrate the validity of the aforementioned argument, this paper also examines the formation pressure characteristics of multiple basins or regions both domestically and internationally. It is found that formation pressure is entirely dependent on the pattern and rate of Cenozoic tectonic movements.

5.1. Example of Active Rapid Overpressure

The Kuqa Depression is a depression with ultra-high pressure, with a maximum pressure coefficient of 2.2. There are numerous research papers on its genesis, mainly suggesting that the intense tectonic compression during the late Himalayan period is the primary factor contributing to the formation of abnormal high pressure in the Kuqa Depression [48,49]. In fact, this region experiences active pressurization. The deposition of an extremely thick gypsum salt layer in the Miocene Jidike Formation (N_1j) slows down the permeation of formation water, followed by the deposition of the Kangcun Formation (N_1K) with a thickness ranging from 200 to 1300 m. Most importantly, the Pliocene Kuqa Formation (N_2K) with a thickness of 450 to 3600 m and the Quaternary Q with a maximum thickness of 1000 m were deposited only 5.2 Ma. Both Dina 202 and Tuzi 2 in the eastern Kuqa Depression are subject to tectonic compression. The former experienced rapid deposition of N_1K + N_2K greater than 3 km, resulting in a pressure coefficient of 2.2, while the latter experienced deposition greater than 1 km, leading to a pressure coefficient of only 1.8 [50]. The pressure coefficient is controlled by the thickness of rapid deposition. This is a typical case of active overpressure.

An abnormal high pressure was observed at approximately 3300 m in the third member of the Shahejie Formation (E_s3) in the Maxi trough, located in the northeastern part of the Raoyang sag in the Jizhong depression. The maximum pressure coefficient reached 1.37 [51]. Similar to the aforementioned Nanpu Sag, the combined thickness of the Guantao Formation, Minghuazhen Formation, and Pingyuan Formation in this region exceeds 2000 m, indicating active overpressure resulting from recent rapid sedimentation and subsidence.

If there is continuous rapid deposition combined with multiple well-sealed gypsum salt layers, the pressure coefficient will be even higher. In the Shizigou area of western Qaidam Basin, there is continuous rapid deposition and subsidence with a total thickness of 2530 m, consisting of 730 m in the Miocene, 1200 m in the Pliocene, and 300 m in the Quaternary Qigequan Formation. At well S23, which is located in the center of gypsum salt deposition, multiple gypsum salt layers are developed, with a single layer thickness of up to 164 m and a cumulative thickness of 325 m. The maximum pressure coefficient in the lower Ganchaigou Formation E₃ (buried depth of about 3800 m) is 2.0, indicating ultra-high pressure [52].

The main source rocks in the Yinggehai-Qiongdongnan Basin are the Oligocene Yacheng Formation and the Miocene Meishan and Sanya Formations, making it the basin with the strongest overpressure among the Cenozoic basins in eastern China and offshore areas. Overpressure is commonly developed below 2800–3000 m, and the measured pressure coefficient (the ratio of formation pressure to hydrostatic pressure) in some intervals is higher than 2.0 [53]. The fundamental reason is the rapid accumulation of sediment during the Quaternary, with a maximum thickness of nearly 2512 m and an extremely high deposition rate.

5.2. Example of Long-Term Static Low Pressure

The Alberta Basin in Canada is a well-studied basin characterized by abnormally low pressure [54,55,56]. The Mississippian to Cretaceous hydrostratigraphic sequence in the southwest region of the basin can be divided into the following two major systems: (1) the lower system consisting of the Mannille Group from the Mississippian to the Lower Cretaceous, and (2) the upper system consisting of the Colorado Group in the upper part of the Lower Cretaceous to the Edmonton Group at the top of the Upper Cretaceous. Since the end of the Cretaceous, the basin has been in a long-term static state, which is consistent with the long-term static low-pressure model.

The main sedimentary strata in the Baise Basin of China include the Eocene Nadu Formation (E_2n), Baigang Formation (E_2b), and the Oligocene Fuping Formation (E_3f) and Jianduling Formation (E_3j). The exposed stratum on the surface is the Jianduling Formation (E_3j), and the Nadu Formation beneath the regional caprock forms a low-pressure system with a pressure coefficient of less than 0.9. In the late Oligocene, the basin began to uplift as a whole, with a stratum denudation thickness greater than 1200 m at the basin margin and greater than 800 m at the basin center [57]. The static time is greater than 23.5 Ma, which is also consistent with the long-term static low-pressure model.

The Songliao Basin experienced folding and atrophy in the Late Cretaceous, undergoing passive uplift and subsidence under compression, followed by a slow cessation of tectonic activity. In the Cenozoic era, it remained in a tectonically inactive static stage for a long time. The eastern Nenjiang Formation (k_2n) was denuded, exhibiting low to normal pressure; only the Qijia-Gulong Depression has 130 m of Cenozoic sediment, showing normal to slightly high pressure.

5.3. Example of Passive Pressurization Mode

There exists passive overpressure in the Yichang area of the Middle Yangtze region. This area has experienced multiple episodes of tectonic uplift and subsidence. Large-scale uplift began at the end of the Triassic period (200 Ma), followed by rapid subsidence in the Early Cretaceous and another rapid uplift during the Miocene and Pliocene (10–5 Ma) [58,59]. This has led to the formation of passive overpressure, with an overpressure coefficient reaching 1.4 times [60]. The theoretical maximum uplift height is calculated to be 1100 m (using the depth of 2800 m in the E’yi Ye-2 well), which is consistent with the results obtained by Tang et al. (2011) [61].

Passive overpressure is also present in the Wuwei Depression of the Lower Yangtze region. In 2019, two sets of abnormally high-pressure natural gas-bearing layers were discovered for the first time in the Triassic Zhouchongcun Formation T_2z of the Wanwei Ye-1 well, with a maximum formation pressure coefficient of 1.9 [62]. The shallow first gypsum salt layer, which is nearly 200 m thick (1975–2175 m), blocks the escape of formation water and natural gas, with a pressure coefficient of 1.79. The second gypsum salt layer beneath it is 50 m thick (2295–2345 m) and has a pressure coefficient of 1.95 (Figure 7). There has been relatively recent rapid uplift and denudation in this area (compared to the Miocene and Pliocene), resulting in slow release of formation pressure within better or thicker cap rocks. Large-scale vertical movement only occurred in eastern China during the Oligocene, with denudation of the Paleogene in southern Jiangsu reaching 2000–3000 m [63]. Zeng (2005) [64] used paleotemperature parameters to reconstruct the tectonic evolution of the Wuwei Depression in the Lower Yangtze region, indicating large-scale uplift and denudation in the study area, resulting in the widespread absence of Jurassic strata, with denudation reaching up to 2550 m during this period. As seen from the aforementioned examples, the pressure release due to uplift and denudation becomes slower as it approaches the present day. Ideally, it should be controlled within 5 Ma, as in the case of uplift in the Sichuan Basin mentioned earlier. The uplift in this area occurred during the Oligocene (37–23.5 Ma). If it were conventional mudstone, the internal high pressure should have been fully released. However, the presence of two gypsum salt layers in the Wuwei Depression has slowed down this process.

Abnormally high pressures have been observed in the Permian and Triassic reservoirs of the Mahu Sag in the northwest Junggar Basin [65], with a maximum overpressure coefficient reaching 1.9 times, exhibiting similar characteristics. During the Neogene period, the basin experienced intense tilting, resulting in uplift in the north and subsidence in the south. The strong subsidence in front of the Northern Tianshan Mountains caused Neogene sediments to accumulate mainly in this area, while the northern part of the basin was predominantly an uplifted and denuded region [66]. Only a portion of the Paleogene strata in the Mahu Sag is preserved in the eastern part. As observed from the northwest–southeast profile across the Mahu Sag, the Cenozoic strata gradually thicken towards the southeast. The minimum residual thickness in the Mahu Sag is 200 m, while it can reach up to 4 km in the Fukang Sag [67]. Based on these observations, it can be inferred that the uplift and denudation in the Mahu Sag are closely related to the intense compression and thrusting from the southern Tianshan Mountains during the Neogene period. The large denudation thickness and relatively recent timing (Pliocene or later) are consistent with the passive pressurization model. Using Formula (2) for this model, the theoretical denudation thickness is calculated to be greater than 4 km.

5.4. Formation Pressure Model

Refer to

Table 4. Overpressure and structural conditions of each basin.

No.	Basin	Pressure Coefficient	Depth of the Top Surface of Abnormal Pressure (m)	Uplift Sediment Thickness (m)	Start Time (Ma)	Lifting and Lowering Rate * (m/Ma)	Literature Source
1	Junggar Basin	1.9	2500	−4000	5.2	−769	Zhang et al., 2022 [65]
2	Qiongdongnan Basin	2.0	2900	2000	2.0	1000	Hao et al., 2015 [26]
3	Qaidam Basin	2.0	2000	1500	5.2	288	Zhang et al., 2016 [52]
4	Jizhong Depression	1.37	3300	2000	3.5	85	Hou et al., 2012 [51]
5	Tarim Basin	2.2	4600	1800	5.2	346	Wang, 2016 [50]
6	Alberta Basin	1.0				0	Wang et al., 2016 [90]
7	Baise Basin	0.9	2000	−800	23.5	−34	Zou et al., 2003 [57]
8	Songliao Basin	1.0	3000	130	83	1.6	Xiang et al., 2006 [41]
9	Lower Yangtze Region	1. 90	2000	−3000	37	−83	Yao et al., 1999 [63]
10	Ordos Basin	0.85	3000	600	23.5	25.5	Li et al., 2012 [47]
11	Sichuan Basin	2.2		−3000	5.0	−600	Deng et al., 2009 [39]
12	Bohai Bay Basin	1.6	3200	2100	5.1	411	Zhang, 2018 [35]
13	West Lake Depression	1.5	3900	8500	24	354	Liu, 2023 [68]
14	Huimin Depression	1.78	3000	1500	25	62.5	Huo, 2020 [69]
15	Erlian Basin	1.0	1300	−180	24	−7.5	Xu, 2007 [43]
16	Pearl River Estuary Basin	1.8	3000	3800	Hanjiang Formation	237	Zhang, 2023 [70]
17	Jiyang Depression	1.7	3400	1500	5	300	Gong, 2005 [71]
18	Kuqa Foreland Basin	2.16	4000	1500	3	179.01	Zhang, 2011 [72]
19	Yinggehai Basin	1.9	4500	500	1.5	250	Mao et al., 2022 [73]
20	Liaohe Basin	1.96	3400	800	8	100	Peng et al., 2013 [74]
21	Yine Basin	1.3	2200	1100	Early Cretaceous	18.33	Zhang et al., 2020 [75]
22	Lunpola Basin	1.0	2100	1000	36.4	25.58	Liu, 2020 [76]
23	Yine Basin	1.429	2000		Early Cretaceous	400	Hou et al., 2019 [77]
24	Hailar Basin	0.81	1700	700	Lower Cretaceous	0	Wang et al., 2010 [78]
25	Yitong Basin	1.1	2500	200	Early-Mid Eocene	190	Wang, 2016 [79]
26	Jiuquan Basin	1.84	3700		Lower Cretaceous	470	Zhou et al., 2013 [80]
27	Yitong Basin	1.0	2304	800		100	Cao et al., 2011 [81]
28	Qaidam Basin	2.1	2300		Eocene Series	330	Han et al., 2023 [82]
29	Gulf of Mexico	1.2	5600	500	29	34.5	Zhang, 2023 [83]
30	Irrawaddy Basin	2.20	1500		Eocene	1500	Chen et al., 2014 [84]
31	Junggar Basin	2.46	2710		Jurassic	500	Lu et al., 2022 [85]
32	Sichuan Basin	1.7		2500	Jurassic	−100	Liu, 2020 [86]
33	Sichuan Basin	1.85	2400		Middle Jurassic	−150	Liu et al., 1995 [87]
34	Middle Yangtze Region	1.4	2700	−1100	8	−137	Zhang et al., 2019 [60]
35	West Kunlun Mountains	2.05	4000	−4000	2	−2000	Ma et al., 2022 [88]
36	Biyang Depression	1.15		300	5	60	Zhang et al., 2020 [89]
37	Fort Worth	1.1	1980	50	24	2	Wang et al., 2016 [90]
38	Guizhou	1.0	1200	80	480	0.1	Wang et al., 2016 [90]
39	Sichuan Basin	1.5		−3100	20.9	−148	Xue et al., 2023 [91]
40	Sichuan Basin	1.18		−700	31	−22.6	Xue et al., 2023 [91]
41	Sichuan Basin	1.1		−1600	26	−61	Xue et al., 2023 [91]
42	Illinois	0.99	−3000	250	64	−3.8	Wang et al., 2016 [90]
43	Qinghai-Tibet Region	1.05	3000	−650	24	−27	Wang et al., 2020 [92]

* Positive numbers represent subsidence, while negative numbers represent uplift.

for more examples. The data in the table are plotted on a semilogarithmic coordinate system of tectonic movement rate v and pressure coefficient c to illustrate their correlation. Positive values represent deposition or subsidence, while negative values represent uplift and erosion. More than 50 famous large basins or regions, both domestically and internationally, were collected. Among them, only 43 basins with both stratigraphic pressure data c and tectonic movement rate v were retained [68,69,70,71,72,73,74,75,76,77,78,79,80,81,82,83,84,85,86,87,88,89,90,91,92]. The data were directly plotted without any filtering, as shown in Figure 8.

Using quadratic curve fitting (dashed line in the figure), the following formula is obtained:

c = 0.13(log₁₀|v|)² + 0.008 log₁₀|v| × sin(v) + 0.95

(4)

where v represents the tectonic movement velocity, measured in meters per million years (m/Ma); c is the pressure coefficient; and sin(v) takes the sin of v. The goodness of fit, R², reaches 0.64, indicating a relatively good fitting effect.

6. Results and Conclusions

Studying the influencing factors of formation pressure with the clue of formation water permeability, the following results are obtained:

(1)

The formation pressure coefficient depends on the permeability of formation water and is positively correlated with the velocity of tectonic movement. The faster the tectonic movement in the Neogene or Quaternary, the more prone to overpressure, while slow tectonic movement in the Cenozoic is dominated by normal or low pressure.

(2)

Compared to pore space, the space vacated by overpressure to normal pressure or the space squeezed is small, less than an order of magnitude observed as abnormally high porosity (5–17%) attributed to undercompaction.

(3)

Physical simulation experiments or numerical simulations of hydrocarbon generation and pressurization are conducted in a completely closed state, accelerating the process of compaction, diagenesis, and hydrocarbon generation faster than the escape rate of formation water, which does not represent real underground geological conditions. Under actual formation conditions, most tectonic activities and hydrocarbon generation processes occur on a time scale that is an order of magnitude larger than the escape rate of formation water. There is no direct correlation between overpressured compartments and the presence of hydrocarbons.

The conclusion is as follows:

(1)

The pressure state of the stratum mainly depends on the mode and intensity of tectonic activity in the late Cenozoic era. Both rapid uplift and rapid subsidence during the late Cenozoic can produce overpressure phenomena, while prolonged slow uplift or neither uplift nor depositional subsidence will result in normal or low pressure.

(2)

There is a positive correlation between the recent tectonic movement’s velocity of ascent and descent and the pressure coefficient.

(3)

Four models of stratum pressure are proposed, i.e., active overpressure model of rapid subsidence (Nanpu Model), passive overpressure model of rapid uplift (Sichuan Model), long-term static normal pressure model (Songliao Model), and large uplift with small subsidence low pressure model (Sulige Model). These models can determine the stratum pressure state within a unified theoretical framework.

(4)

A method for calculating the amount of denudation using the characteristics of formation water pressure relief, as well as a method for calculating the maximum height of tectonic uplift, are proposed.

(5)

Other hypotheses for the causes of overpressure, such as undercompaction, tectonic compression, and hydrocarbon generation pressurization, may exist. However, their impact on overpressure is not significant.

This article will help us gain a new perspective on the concept of overpressure formation mechanisms related to sedimentology and petroleum geology.